Evidence for a plate tectonics debate

Ž .Earth-Science Reviews 55 2001 235–336www.elsevier.comrlocaterearscirev

Evidence for a plate tectonics debateq

MacKenzie Keith)

Department of Geosciences, PennsylÕania State UniÕersity, UniÕersity Park, PA, 16802, USA

Received 17 October 2000; accepted 4 April 2001

Abstract

A central problem in understanding the Earth system is the relationship between mantle convection and near-surfaceŽ .structure, geophysics and geochemistry. Unexplained anomalies of the plate tectonics PT model are the postulated

mechanism of generating the axial rift valley by upward movement of marginal fault blocks, contrary to observations in theIcelandic rift, and the mechanism by which a postulated broad upwelling plume can yield the required narrow zone of axialvolcanism. These discrepancies are the basis for a critical reappraisal of other evidence. It is shown that the generation ofoceanic magnetic stripes, which led to the ingenious spreading hypothesis, is a result of narrowing of a formerly expandedocean ridge volcanic system, and resultant sequential cooling of crust and upper mantle, and does not require ocean floorspreading. An alternative ridge model postulates sub-continent upwelling, sub-ridge convergent flow and generation of

Ž .mid-ocean ridge basalt MORB by heterogeneous volatile-promoted melting at sub-axial zones where oceanic crust isrecycled into the mantle. That model is favored by the results of convection experiments and confirmed by a series of

Ž .independent indicators: heat flow beneath ridge flanks, coriolis curvature of fracture zones FZ , downstream development ofconvective rolls and near-transform tectonic rotations. Sub-axial downflow is confirmed by North Atlantic positive geoidanomalies, by high P-wave velocities deep below the ridge axis, and by the synclinal AflexloadB structure and compressionalstress regime of near-axial crust. Oceanic island volcanism is attributed to the same process, crustal recycling at local sites of

Ž . Ž .downflow focused by deep residual masses that are relatively cold radioactivity depleted and viscous volatile depleted .The recycling model is confirmed, for both MORB and OIB, and decompression melting of mantle plumes rejected, on

Ž .several grounds: 1 local recycling of crust and resultant mantle hydration are indicated by diapirs of serpentinized mantleŽ .along fracture zones and ridge axes; 2 distinctive isotopic signatures in oceanic basalt require local recycling of crust and

Ž .sediments to the magma source; 3 the principal oceanic island groups, exemplified by the Hawaiian islands, lack thepositive heat flow anomaly expected over a plume and are underlain by zones of relatively high seismic velocities. Theproposed local recycling accounts for several previously puzzling features of global geochemical systems, including thecarbon budget, the lead paradox and the isotopic array of heavy noble gases in oceanic basalt. Deduced end-membercomponents, defined by isotope ratios, can be attributed to regionally variable mixtures of upper mantle with terrigenous andpelagic sediments, subaerial and suboceanic basalt, and with depleted lower mantle, proposed to have been available duringa Pacific-centered Mesozoic mantle surge. There is no need to appeal to the several currently favored isolated reservoirs inthe upper and lower mantle. Seismic data for continental ranges, exemplified by the Alpine region, show that the plate

q The editors regret to have learned of the death of MacKenzie Keith shortly after this paper was submitted. In these circumstances theyhave decided to publish the article as submitted.

) Tel.: q1-814-867-9032.Ž .E-mail address: [email protected] M. Keith .

0012-8252r01r$ - see front matter q 2001 Elsevier Science B.V. All rights reserved.Ž .PII: S0012-8252 01 00060-5

( )M. KeithrEarth-Science ReÕiews 55 2001 235–336236

collision model is misleading in view of increasing evidence for upper mantle deformation by viscous creep. Continental riftsystems are re-examined on the basis that the oceanic axial zone can be traced to the African rifts, via Afar, and are thereforetaken to be a result of a similar dynamic system. It is shown that dominant upwelling plumes are focused beneath the thickkeels of ancient cratons and that rift systems, typically at craton margins, are collapse structures in zones of convergence anddownflow. q 2001 Elsevier Science B.V. All rights reserved.

Keywords: plate tectonics; mantle convection; geophysics; geochemistry

Contents

1. Sub-ocean mantle flow and genesis of MORB. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2371.1. Clues to direction of mantle flow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 238

1.1.1. Experimental convection . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2381.1.2. Thermal barriers in the Earth . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2381.1.3. Sub-continent outflow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2391.1.4. Sub-ridge heat flow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2391.1.5. Coriolis curvatures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2401.1.6. Downstream convective rolls . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2401.1.7. Near-transform rotations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2411.1.8. Structure of axial and fracture zone valleys . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2421.1.9. Axial and fracture zone gravity anomalies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2441.1.10. Broad geoid anomalies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2441.1.11. Seismic indicators of sub-ridge convergent flow. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2451.1.12. Sub-axial seismic velocities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2451.1.13. Transition zone deflections . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2481.1.14. Mantle diapirs and uplifted old crust . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2491.1.15. Clues to sub-fracture zone and sub-axial recycling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2511.1.16. Vesicle pressure in MORB glass . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 252

1.2. Oceanic mantle dynamics and surface expressions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2521.2.1. A viscous boundary-layer model of upper mantle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2521.2.2. Convective drive . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2531.2.3. Icelandic structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2531.2.4. Tilted crust and a AflexloadB model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 255

Ž .1.2.5. Rock stress, geodetic and global positioning system GPS data and earthquake mechanisms . . . . . . . . . . . . . . 2561.2.6. Axial and transform earthquakes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2571.2.7. Flexural constraints on volcanism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 258

1.3. Mesozoic mantle surge . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2591.3.1. Disruption of steady-state convection . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2591.3.2. Expansion and elevation of oceanic ridges . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2611.3.3. Widening of the MORB volcanic zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2611.3.4. Limits of surge-related MAR volcanism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2621.3.5. Contrasting features of MAR and EPR . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2651.3.6. Subaerial exposure and subsequent sinking of the ridge . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2661.3.7. Volcanic zone contraction and the generation of magnetic stripes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2681.3.8. Ages of sediment caps and interlayers on the ridge . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 271

2. MORB recycling, cold spots and oceanic islands . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2732.1. A recycling model for genesis of MORB . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 274

2.1.1. Oceanic LVZ and the MORB source . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2742.1.2. Volatile components of MORB . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2752.1.3. Volatile-promoted melting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2762.1.4. Local sub-axial recycling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2782.1.5. Sites of sediment accumulation and recycling. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2802.1.6. Sedimentary source of oceanic sulfide deposit Pb. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 281

( )M. KeithrEarth-Science ReÕiews 55 2001 235–336 237

2.2. Global geochemical systems . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2822.2.1. Steady-state and episodic recycling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2822.2.2. The carbon budget . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2822.2.3. The lead paradox . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2832.2.4. Heavy noble gases in oceanic basalt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 283

2.3. Reservoirs and recycling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2852.3.1. Geochemical constraints on the deep-plume model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2852.3.2. Principal reservoirs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2852.3.3. Episodic lower mantle contribution . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2872.3.4. Sediment signatures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2882.3.5. Examples of source mixing . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2892.3.6. Osmium isotope ratios . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2902.3.7. Mesozoic submarine basalts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2912.3.8. Subaerial MORB and cosmogenic helium-3 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 292

2.4. Cold spots and oceanic islands . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2922.4.1. Geophysical clues to ocean island dynamics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2932.4.2. Sub-island recycling and melting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2952.4.3. Sources of high 3Her4 He ratios . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2972.4.4. The helium–neon association . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 299

2.5. Pacific disruption and oceanic island chains . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 301

3. Continental rifts and ranges. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3043.1. Continental mountain ranges . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 304

3.1.1. SKS data and interpretations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3053.1.2. Sub-alpine anisotropy and deduced mantle flow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3073.1.3. Application to other mountain ranges . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 309

3.2. Continental rifts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3093.2.1. The rift connection: ocean to continent . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3093.2.2. Proposed rift model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3103.2.3. Structural and geophysical evidence . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 310

3.3. Sub-rift recycling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3123.3.1. Diversity of rift volcanism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3123.3.2. Helium isotopic variation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3133.3.3. Genesis of rift carbonatites. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3133.3.4. Carbon dioxide emissions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3143.3.5. Oldoinyo Lengai, key to the rift system . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 315

4. Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 318

Acknowledgements. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 318

References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 318

1. Sub-ocean mantle flow and genesis of MORB

Ž .Essential features of the plate tectonics PT hy-pothesis are widely accepted but some aspects of themodel are open to question because they are inapparent conflict with known properties of materialsand of the Earth system. The first question applies to

Žthe conventional rigid plate hypothesis Turcotte and.Schubert, 1982; Morgan, 1998 , as restated in a

Ž .recent review Kearey and Vine, 1990, p. 73 :

AWithin the basic theory of plate tectonics, plates areconsidered to be internally rigid, and to act as ex-tremely efficient stress guides. A stress applied toone margin is transmitted to its opposite margin withno deformation of the plate interior.B That concept isclearly inconsistent with the results of experiments

Žon rock strength Handy, 1990; Kohlstedt et al.,.1995 and with the scale factors that govern the

strength of large masses of rock on a geological timeŽ .frame Hubbert, 1937 , both of which lead us to


imply that large rock masses are weak and subject tohot creep. The stability of some hydrous phases at

Župper mantle conditions Schmidt and Poli, 1994;.Bose et al., 1996 , and the identification of volatiles

Žin micro-inclusions in diamond Schrauder and.Navon, 1994; Leung et al., 1994; Haggerty, 1997 ,

are taken to indicate hydrated upper mantle, andorder of magnitude reductions of viscosity and creepstrength relative to anhydrous olivine rheologyŽChopra and Paterson, 1984; Hirth and Kohlstedt,

.1996 .There is also an evident need to re-examine the

hypothesis of sub-axial upwelling. How can the broadŽplume of passive flow models MELT seismic team,

.1998 generate the narrow zone of axial volcanism,the Aknife-edgeB separation of adjacent flow regimeson either side of a transform, and the unbelievable

ŽAoverlapping spreading centersB Macdonald et al.,.1986b, 1988; Barth, 1994; Cormier et al., 1996 ?

Further questions are related to the concept of oceanfloor spreading and the Atape recorderB model forgenerating the oceanic magnetic stripes. The spread-ing model has an appealing simplicity, serves toaccount for the magnetic stripes and is widely ac-cepted, despite several puzzling discrepancies, in-cluding the occurrence of non-matching geochemical

Žanomalies on opposite ridge flanks Langmuir et al.,.1992 , the requirement for uplift of marginal blocks

Žto form the axial valley walls contrary to evidence.in Iceland , and the failure to find the required

narrow zone of crustal accretion with the property ofdividing neatly so that matching halves move toeither side.

Evidence is summarized below to show that nospreading is required. The oceanic magnetic stripescan be accounted for by narrowing of the mid-oc-eanic volcanic zone and consequent crestward migra-tion of the blocking temperature following a Meso-zoic surge of mantle flow and volcanic zone expan-

Ž .sion that peaked in Cretaceous time Section 1.3.7 .ŽAn alternative model of upper mantle flow Keith,

.1993 postulates sub-continent upwelling, sub-conti-nent to sub-ocean outflow at Atlantic-type margins,and convergent sub-ocean mantle flow toward theaxes of mid-ocean ridges, defined as principal con-vection cell boundaries. A wide array of criticalevidence is summarized below, in simplified formŽ .Section 1.1 , reduced to a key question regarding

the direction of sub-ocean mantle flow, i.e. divergentŽ .flow plate tectonics model or sub-ridge conver-Ž .gence alternative model . The important clues in-

volve several fields of the geosciences and thereforeare presented, as much as possible, in a form suitablefor the general reader, leaving out much of thebackground and analysis that might be expected by aspecialist, for example in structural geology, petrol-ogy, geochemistry or seismology. The weight ofevidence clearly supports the alternative model andis contrary to the spreading model of plate tectonics.

1.1. Clues to direction of mantle flow

1.1.1. Experimental conÕectionExperiments show that typical upwelling zones

are broad plumes and that the only linear features onthe upper surface of a convecting fluid are the tracesof narrow planar zones of convergence and down-flow at cell boundaries. Numerical simulations of

Žmantle convection McKenzie, 1979; Bercovici et.al., 1989a,b; Christensen and Harder, 1991 are con-

sistent with the convection experiments; they showthat upwelling has the form of broad cylindricalplumes, and that the typical downwelling mode isdominated by relatively thin planar sheets. TackleyŽ .1996 showed that other geometrical modes arepossible, but that upwelling plumes and downwellingsheets are favored under any one of several condi-

Ž .tions that are relevant to mantle convection: 1 aŽ .contribution from internal heating; 2 depth-depen-

Ž .dent compressibility and associated properties; 3Ž .spherical geometry; and 4 temperature-dependent

viscosity.A consistent feature of experimental convection

systems, whether heated from within or below, isthat a thermal barrier in the lid creates a local hotspot and lateral gradients in the upper boundarylayer, with the result that each barrier focuses an

Ž .upwelling plume Keith, 1993 . In that type of sys-tem the size and number of convection cells isdetermined by the dominant thermal barriers and thesurface pattern is characterized by cell boundariesmidway between them.

1.1.2. Thermal barriers in the EarthGeophysical data are consistent with the results of

convection experiments; they indicate that the domi-


nant thermal barriers in the Earth system are ancientlithospheric keels. The continental keels are typicallyabout 150–250 km thick below the cratonic cores of

ŽArchean shield areas Boyd and Gurney, 1986; Keith,.1993, 1994; Pearson et al., 1995a,b . They are effec-

tively isolated from convection by their buoyancyand by high viscosity, presumably due to Archean

Žpartial melting events and loss of volatiles Doin et.al., 1996 . They will function as effective thermal

barriers above the convecting mantle, and it followsthat dominant upwelling plumes are to be expectedbeneath continental shields, not beneath mid-oceanridges. That association is consistent with the nega-tive geoid anomalies over major shields, and withindications of relatively high temperatures in sub-

Ž .shield mantle, including a large q thermal anomalyŽin sub-African lower mantle Dziewonski et al., 1993;

.Ritsema et al., 1998 . Recent interpretations of seis-mic tomography, based on P- and S-wave data, ledto the conclusion that the uplift of southern andeastern African plateaus is a dynamic effect related

Žto sub-craton upwelling Grand et al., 1997; Lith-.gow-Bertelloni and Silver, 1998 .

Further evidence of deep cratonic keels is pro-vided by the occurrence of diamonds in cratonic

Ž .kimberlites Haggerty, 1986, 1997 , and by kimber-lite-rafted mantle xenoliths, which yield a petro-graphically defined geotherm inflection, below theAfrican Kaapvaal craton, for example, and a corre-

Žsponding change from coarse granular texture the.continental keel to finer-grained sheared textures

Ž .asthenosphere in the depth range 175–250 kmŽ .Nixon, 1987; Keith, 1993 . That change is taken torepresent the top of a viscous boundary layer cap-

Žping the asthenosphere see Anderson; Boyd and.Mertzman, in Mysen, 1987 .

Ž .Ballard and Pollack 1987 noted the contrastbetween low cratonic heat flow, typically about 40

2 Ž 2 .mWrm , and the higher heat flow 65–70 mWrmin surrounding Proterozoic and younger mobile belts.Their model, based on data for the Kaapvaal andZimbabwe cratons and surroundings, made al-lowance for differences of lithospheric heat produc-

w xtion. They concluded op. cit., p. 256 that convec-tive diversion of warm upwelling mantle away fromthe base of the craton and into the base of surround-ing mobile belts provides a single dominant mecha-nism that regulates the heat flow of the region.

Ž .Nyblade et al. 1990 reported a similar heat flowpattern around the Tanzania craton of equatorial eastAfrica, and later concluded that the pattern is typical

Žof ancient cratons and their surroundings Nyblade.and Pollack, 1993 .

1.1.3. Sub-continent outflowThe evidence for sub-craton upwelling leads to

the question as to whether the extensional tectonicsŽof Atlantic-type continental margins Tankard and

.Balkwill, 1989 are due to outward sub-continentmantle flow. That possibility is consistent with themarginal gravity anomalies, best accounted for by asteady outward creep of low-density continental crustŽ .Worzel, 1965; Bott, 1971, p. 90 . The proposedoutflow model gains further support from the con-straints of global heat flow; sub-oceanic heat lossfrom outward flowing oceanic mantle serves to ac-count for a long-standing puzzle: the equality ofcontinental and oceanic heat flow. An essential fea-ture of the global system, consistent with the bal-anced heat flow, is that mantle convection is aresponsive, self-regulating system.

1.1.4. Sub-ridge heat flowThe oceanic heat flow data are subject to different

interpretations, dependent on whether allowance ismade for the effect of anomalous heat flow measure-ments near centers of volcanism or hydrothermalactivity. Widespread acceptance of the view thatvolcanic activity is restricted to a narrow axial zoneŽ .plate tectonics model has led investigators to con-struct heat flow profiles that include high and widelyvariable data from the near-crestal zone and conse-quently to deduce a crestward increase in averageheat flow. It became evident, however, from early

Ž .studies of heat flow variability Lister, 1972 thathydrothermal circulation is the dominant heat trans-fer process in the crestal zone, even at the rare sitewhere there is a thick sediment cover, as on Juan de

Ž .Fuca Ridge Baker et al., 1998; Stein et al., 1998 .Ž .Indications of off-axis volcanism Macdonald, 1986

and a wide zone of hydrothermal circulation, out toŽcrust dated at about 65 Ma Mottl and Wheat, 1994;

.Booij and Staudigel, 1996 , lead one to question thesignificance of near-crestal heat flow data. One canminimize the effects of local volcanic distortions andemphasize the regional trend of conductive heat from


the mantle by focusing on data from outer flanks ofthe ridge, outside of the zone of active volcanismand where hydrothermal seawater circulation is sealedoff by continuous thick sediment cover. The signifi-cant result is that North Atlantic ridge flank datashow a progressive decrease in heat flow toward the

Ž .ridge axis Fig. 1 , consistent with convergent flowof oceanic upper mantle. Indian Ocean Ridge heatflow data show a similar decrease toward the ridge

Ž .crest Keith, 1993 .

1.1.5. Coriolis curÕaturesThe Coriolis deflection applies in varying degrees

to all flow on the Earth, and increases with latitude.As a consequence of Earth rotation, all flow pathsveer to the right in the northern hemisphere, to theleft in the southern hemisphere. If the effect isappreciable for slow rates of mantle flow, it shouldbe evident in major lineations such as oceanic frac-

Ž .ture zones FZ . On that basis, one cannot easilydismiss the observation that geoid lineations associ-ated with the longest high-latitude fracture zones,Eltanin and Udintsev in the south Pacific, Charlie–Gibbs in the North Atlantic, have regional curvatures

consistent with convergent flow of upper mantleŽ Ž .toward the ridge axis see Keith 1993 and charts of

Ž . Ž ..Haxby 1985 and Smith and Sandwell 1997 .

1.1.6. Downstream conÕectiÕe rollsLongitudinal convective rolls are judged to be the

principal form of secondary convection within theupper boundary layer of sub-ridge mantle, a low-viscosity region estimated at 75–125 km thick, basedon initial 150–250 km wavelength of linear geoid

Ž .anomalies Robinson et al., 1988 . It is surely notaccidental that the estimated thickness of a low-viscosity zone matches that of the seismic low-veloc-

Ž .ity zone LVZ . The development of rolls in theupper mantle was suggested some years ago, based

Ž .on a numerical model Richter, 1973 and on priorŽ .convection experiments Keith, 1972, 1993 . A sub-

Žsequent experimental study Richter and Parsons,.1975 served to confirm that longitudinal rolls are

the preferred mode of secondary convection underconditions of lateral flow and top cooling. A compar-ison of experimental rolls with fracture zones on the

Ž .Mid-Atlantic Ridge MAR led to the proposal thatfracture zones and transforms are surface expressions

Ž .Fig. 1. Heat flow versus magnetic anomaly age for the North Atlantic, redrawn from Sclater and Wixon 1986 to show an arbitrary divisionŽ . Ž .into two populations: 1 widely variable data from the crestal and near-axial zone open circles , notably distorted by hydrothermal

Ž . Ž .circulation; and 2 less variable data from outer flanks and marginal basins filled circles . The filled circles represent high-quality heatŽflow data for the region of continuous sediment cover beyond the 0 to 60 Ma zone of active seawater circulation through the crust Mottl

.and Wheat, 1994 .


Žof roll boundaries and are zones of downflow Keith,.1972, 1993 . That proposal is consistent with the

typical occurrence of serpentinized mantle diapirsalong fracture zones, an indicator of recycling ofhydrous crustal materials. The fracture zone:cellboundary association is further supported by theobservation that Pacific geoid profiles show a narrownegative anomaly over a major fracture zoneŽ .Udintsev , taken to imply sub-fracture zone down-

Ž .flow Robinson et al., 1988, fig. 17 and p. 34 .Several investigators of convective rolls have pro-

posed that rolls are expressed as linear gravityanomalies but did not make the fracture zone con-nection, possibly because of an observed angulardifference, in the Pacific, between the trend lines of

Žfracture zones and gravity anomalies Haxby and.Weissel, 1986; Baudry and Kroenke, 1991 . That

angular difference can be understood if some of thePacific fracture zones are residual from an earlier

Žregime of mantle flow, centered in the Pacific seeSection 3 about proposed Mesozoic surge and post-

.surge return toward a steady-state system .A significant observation from the convection

experiments is the evidence that convective rolls arenot observed near an upwelling zone; they are atypical downstream feature of a system subject tosurface cooling. The convection experiments of

Ž .Curlet, illustrated by Parsons and McKenzie 1978 ,are particularly relevant. Lateral flow in Curlet’stank experiment was induced by uniform top heatingof one side and uniform top cooling of the other side,a configuration that might be taken as a simplemodel of outflow from beneath a continent andsubsequent heat loss in a surrounding oceanic region.A notable result, in the downstream region of lateralflow and heat loss, was the development of an upperboundary layer that eventually became unstable andyielded a regular pattern of convective rolls alignedin the direction of principal flow. A similar patternof downstream convective rolls is obtained by substi-tuting a thermal barrier in the lid, in place of thesuperimposed heat source, or by heating one part of

Žthe experimental fluid from within or below Keith,.1993, p. 208 . Any AspokeB pattern of Rayleigh–Be-

Ž .nard convection cells Busse and Whitehead, 1971evolves into elongated rolls under the influence of

Ž .dominant lateral flow. Parsons and McKenzie 1978calculated, for assumed viscosity of 1021 Pa s, that

the upper mantle thermal boundary layer could gounstable and develop rolls only after a minimal 70Ma of cooling. It follows that the exclusive occur-rence of rolls in the young near-axial region ofmid-ocean ridges, identified either as a fracture zone

Ž .pattern North Atlantic or as a short-wavelengthŽ . Žlinear geoid anomaly pattern eastern Pacific Haxby

.and Weissel, 1986; Maia and Diament, 1991 , identi-fies the axial region as the downstream end ofsub-ridge lateral flow. The evidence is consistentwith the proposed sub-ocean convergent flow andcontrary to the plate tectonics model of mid-oceanupwelling and divergence.

1.1.7. Near-transform rotationsA simple structural indicator of the direction of

mantle flow is the observed curvature of ridge mor-phology and structure elements adjacent to majortransforms. A prime example from Atlantis FZŽ .Blackman et al., 1998 is the basis for Fig. 2. Other

Žexamples include the Charlie–Gibbs FZ Searle,. Ž .1979 , Vema FZ Macdonald et al., 1986a , Clipper-

Žton FZ in the Pacific Gallo et al., 1986; Macdonald. Ž .et al., 1988 , and Kane FZ Kong et al., 1988 . The

observed rotations have conventionally been at-tributed to faulting and dike injection within a ro-

Žtated stress field Searle, 1983; Macdonald et al.,1986a; Fox and Gallo, 1986, p. 168; Gudmundsson,

.1995 . They are proposed, instead, to be topographicrotations, an expression of persistent traction at aboundary between the flow regimes of oppositelymoving convective rolls. That proposal is favored for

Ž .three principal reasons. 1 The dominant near-surface transform valley faults are not the Riedelshears of the rotated stress field model; they are

Ž . Ž .dip–slip faults Fox and Gallo, 1986, p. 161 . 2Near-transform rotations are not restricted to faultsand dikes; they include all features of ridge and

Ž .valley topography. 3 A recent study at ridge–frac-Ž .ture zone intersections RFZ in the North Atlantic

Ž .Tanaka et al., 1996 showed that some magneticlineations are curved near the fracture zone, mimick-ing the topographic rotations. In view of the evi-dence that oceanic magnetic stripes are due to late,

Ž .post-tilting magnetization Cande and Kent, 1985 ,the curved magnetic stripes provide clear evidencefor post-magnetization tectonic rotation rather thanmagmatic intrusion into a distorted stress field. The


Ž .Fig. 2. A Typical structure and topography of a North Atlantic transform valley and of intersections with the axial rift valley of theŽ . Ž .Mid-Atlantic Ridge redrawn from Fig. 1b of Fox and Gallo, 1986 , combined with the Atlantis transform chart of Blackman et al. 1998 ,

to show tectonic rotations adjacent to a typical fracture zone and transform. Isolated uplift hills have been identified as massifs ofŽserpentinized peridotite, interpreted as upper mantle diapirs Bonatti and Honnorez, 1976; Auzende et al., 1994; Charlou et al., 1997;

.Blackman et al., 1998 . Some mantle diapirs are capped with basaltic crust. Others are unroofed and near-transform exposures showŽ .transform-parallel linear structure proposed to be shear-zone mylonite in the upper mantle see text . Nodal basins at ridge–transform

Ž .intersections are interpreted as zones of strongly focused downflow. B Interpretation of diagram A, to show proposed axial downwellingŽ .zones two heavy lines and parallel line shading . Arrows represent deduced flow directions and relative flow rates of sub-ridge mantle,

Ž .based on the proposed model and on the postulate that fracture zones are developed at the boundaries of convective rolls see text , and thatŽ .the observed physiographic rotations at their margins are not due to magmatic intrusion into a distorted stress field plate tectonics model

but instead are physical rotations of all crustal structures, the result of the traction of differential mantle flow in a convergent flow regime.Differential flow rates beyond the transform limits are attributed to progressive reduction of the horizontal component of flow as the axialdownflow region is approached.

observed near-transform and near-fracture zone cur-vatures therefore indicate convergent mantle flow

Ž .toward the ridge axis Fig. 2 .

1.1.8. Structure of axial and fracture zone ÕalleysAxial valley structure and similar structures of

fracture zones provide key evidence regarding mid-ocean ridge dynamics. Axial valley relief on the

ŽMAR is in the range 600–2000 m average ;1300.m , and the valley widths range from 16 to 62 km

Ž .average 35 km . Neither the relief nor the width areŽ .correlated with Aspreading rateB Malinvemo, 1988 .

The structures of Atlantic valley-fill sediments arenot clearly defined in typical seismic reflection pro-files, presumably because of minimal supply of clas-tic sediments to the crestal zone and the dominanceof uniform, fine-grained pelagics. There are indica-tions that the valleys are filled with rubble from thevalley walls, including fragments of basalt, gabbroand serpentinized peridotite. A seismic refraction


Ž .survey along the Oceangrapher FZ Fox et al., 1976 ,combined with measurement of the acoustic proper-ties of locally dredged rock specimens, led to theconclusion that the fracture zone valley is underlain

Ž .by a 2-km-thick layer V s4.4 kmrs deduced toP

be basaltic rubble or serpentinite, and a lower layerŽ .V s6.5 kmrs best approximated by serpentinizedP

ultramafic rock.More clearly defined layering is evident in parts

Ž .of the East Pacific Rise EPR that are close to thecontinent and thus receive an abundant supply ofvariable detrital sediment. Seismic reflection surveysacross the well-sedimented median valleys com-monly show synform deformation of the sedimentaryand volcaniclastic layers and a downward increase inthe synclinal sag, the expected form of a collapse

Žstructure and of downward-moving valley fill Davis.and Lister, 1977 . One of the Juan de Fuca profiles

Ž .their fig. 7a shows compressional folding of theŽ .valley fill. Another their fig. 6 shows tilted and

synform lower beds unconformably overlain byundistorted layers of the youngest sediments. Re-markably similar sag structures, consistent with asubsidence regime, are evident in seismic reflectionprofiles across well-sedimented fracture zone val-

Žleys, for example across the Kane FZ Jaroslow and.Tucholke, 1994 , the Owen FZ in the Indian Ocean

Ž .Bonatti et al., 1979 and across the deep nodalŽ .basins at typical ridge–transform intersections RTI

Ž .Rowlett and Forsyth, 1984 .The typical axial valley is dominated by dip slip

Ž .faults Fig. 2 as indicated by observations, fromsubmersible, of down–dip striations and slickensides

Ž .on fault surfaces Fox and Gallo, 1986 , and by thenormal fault mechanisms of axial zone earthquakes.The dominance of dip–slip faults in the axial valleyis consistent with the collapse model of rift valley

Ž .formation Keith, 1993 and with the original defini-Žtions of rift and graben Gary et al., 1965, Glossary

.of Geology . On the other hand, the normal faultcharacter of the valleys has been adapted to the platetectonics model by the assumption that new oceancrust and lithosphere are generated in a narrow cen-tral zone and are subsequently uplifted in a sequence

Žof fault blocks that form the valley walls Ballardand Van Andel, 1977; Macdonald, 1986, p. 58;

. ŽMcAllister and Cann, 1996 . Vogt Vogt and Tu-.cholke, 1986, p. 419 described the hypothetical

mechanism as an upward-moving Atectonic escala-torB. There is an apparent need to re-examine theinterpretations of axial zone earthquakes, on thegrounds that upward moving blocks may yield sourcemechanisms different from those of collapse struc-tures, and possibly similar mechanisms of landslides.

Ž .A recent study Miller et al., 1998 provided evi-dence that some types of earthquakes, such as thoseexcited by landslides and volcanic eruptions, areconsistent with net forces rather than the conven-tional double couple mechanism.

The question of block uplift versus block sinkingcan be subjected to simple tests in Iceland, or in therift system in East Africa, where valley subsidence is

Ž .evident Holmes, 1965, pp. 1059–1078 . The Ice-landic evidence is particularly clear-cut and unam-biguous. Where the axial zone extends from theMAR to Reykjanes Peninsula, a classic rift valley is

Ž .developed at Thingvellir T of Fig. 5 . One can walk,as I have done, on the marginal plateau above the riftvalley, follow a specific boundary between an olderand a younger lava flow and trace it to the faultscarp at the valley margin. From that cliff-top van-tage point, one can look down and see the continua-tion of the selected lava flow boundary on subsidedfault blocks below and on tilted blocks mixed with

Žrubble at the base of the cliff cf. photo in Matthews,.1973 . It is evident that the marginal blocks are

moving downward relative to the bordering plateau.One can also look backward, away from the rim, andnote the absence of clearly defined fault blocks onthe marginal plateau. It is appropriate to ask how the

Ž .uplifted blocks plate tectonics model can moveupward from the rift floor to form the valley wallsand then be so neatly relevelled and fitted togetherthat their boundaries disappear.

wNothing is as sobering as an outcrop attributed toxFrancis Pettijohn

Ž .The indications above that the oceanic rift valleyis a downward-moving collapse structure leads to are-examination of the rift-parallel open fissures thatoccur in places on the rift valley floor. Proponents ofplate tectonics have taken the gaping Icelandic fis-sures as evidence of rift-normal extension andspreading but the typical fissures develop along the

Žcrests of elongated bulges see example in Matthews,.1973 . Similar fissured bulges, referred to as axial


Ž .volcanic ridges AVRs or neovolcanic ridges, havebeen mapped in the valley floors of many sub-ocean

Žsegments of the axial valley Ballard and Van Andel,1977; Macdonald, 1982; Sempere et al., 1993; Head

.et al., 1996 . A prime example is shown in thethree-dimensional perspective view of Smith et al.Ž .1997 . The AVRs are typically 100–500 m high,1–3 km wide, and up to 40 km long. They typicallyhave gently sloping sides and summit fissures andare the principal sites of recent fissure eruptions,leading one to deduce that they are bulges producedby shallow level magma injection, probably at the

Ž . Ž .depth of level of neutral buoyancy LNB Fig. 6 .The orientation of the fissured bulges, parallel to therift valley, provides evidence of universal rift-normalcompression, as measured in the exposed Icelandic

Ž .sector Section 1.2.5 , and favors the proposal thatrift valleys are not indicators of spreading; they arecollapse structures within a compressional regime.

Remarkably, in view of the postulated differentŽorigins of axial and transform valleys plate tectonics

.model , and the occurrence of strike–slip earth-quakes at mid-crustal depths on transforms, thetransform valleys exhibit dominant dip–slip faults

Ž .similar to axial valley faults Fig. 2 , and the valleyshave similar dimensions, with widths in the range

Ž .30"10 km Macdonald, 1986; Dick et al., 1991 .The dip–slip character of dominant faults in bothvalley types, strongly suggestive of a collapse struc-ture, is accentuated in the nodal basins at RTI in the

ŽAtlantic and Indian Oceans cf. Macdonald et al.,.1986a, p. 3351 .

In the proposed model, the above similarities ofaxial and transform valleys are to be expected; bothare collapse structures at sites of downflow. Fracturezones are taken to be the surface expression ofshallow subduction at the boundaries of convectiverolls, and the nodal deeps at RTI are to be anexpression of strongly focused downflow at the con-fluence of roll margin boundaries with the deeperdownflow regime below the axial zone.

1.1.9. Axial and fracture zone graÕity anomaliesIn Atlantic ridge segments with a well-developed

axial valley, some Bouguer anomaly profiles show aŽnarrow negative anomaly over the axial valley Pariso

.et al., 1995; Fujimoto et al., 1996 , presumablyŽ .indicating low-density valley fill. The axial y

Ž .anomaly typically is bordered by a broad qanomaly. The investigators suggested a plate tecton-

Ž .ics hypothesis plate cooling to account for theoutward change from a negative to a positive Bougueranomaly, but the plate cooling hypothesis cannoteasily account for the narrow dimensions of the axialnegative anomaly. A broad negative anomaly wouldbe expected above an upwelling mantle plume.

Ž .Neumann and Forsyth 1993 investigated theBouguer gravity anomalies and bathymetry of theMAR axial region and showed that the median val-ley topography is not compensated by crustal thick-ness or density changes, and therefore must be cre-ated and maintained by a dynamic process. They donot suggest such a process but the simplest one issub-axial downflow with resultant downward trac-tion on a subrift wedge of crustal material.

The gravitational anomalies over fracture zonesand transforms differ from the typical axial anoma-lies, a difference that can be attributed, in part, to themore common occurrence of serpentinite and serpen-tinized peridotite along the fracture zones, presum-ably a result of relatively shallow recycling of hy-drous crustal material. Seismic investigations ofAtlantic fracture zones led to the conclusion thatthey are characterized by relatively thin crust over

Ž .updomed serpentinized mantle Detrick et al., 1993 .The diapirs of serpentinized mantle peridotite formextensive transverse ridges along some fracturezones, for example along the Vema and Romanche

Žfracture zones in the equatorial Atlantic Bonatti and.Honnorez, 1976; Auzende et al., 1989 , but more

commonly generate discrete massifs, preferentially atŽthe inside corners of the RTI Juteau et al., 1990;

.Blackman et al., 1998 . Examples are shown in Fig.2.

1.1.10. Broad geoid anomaliesThe current assumption that mid-ocean ridges are

sites of upwelling has led to a persistent puzzle,Ž .clearly stated, e.g. by Koch and Ribe 1989, p. 538 :

AAn unresolved problem in geoid modeling is theassociation of geoid highs with both sinking slabsand rising plumesB. Recent summaries of geoid

Žanomalies Richards et al., 1988; Heirtzler and Fraw-. Ž .ley, 1994 show that after removal of the strong q

gravitational contributions from Indonesian and Pa-cific subduction zones, the main positive residuals


are associated with mountain ranges, including someŽ .mid-ocean ridges, notably the major q anomaly

over the northern MAR, a long established ridgesegment that was relatively protected from the Pa-

Ž .cific-centered Mesozoic disturbance Keith, 1993 .Some investigators, in attempting to reconcile the

Ž .observed mid-ocean ridge q anomalies with pre-sumed sub-ridge upwelling, have assumed that thepositive effect of bulging of the surface and ofinternal interfaces overwhelms the negative effect oflow density in a rising plume, thus yielding a net

Žpositive geoid anomaly McKenzie et al., 1973; Tan-.imoto and Anderson, 1984, p. 287 . That interpreta-

Ž .tion has been questioned Richards et al., 1988 , andthe total contribution of surface topography andnear-surface density contrasts has been estimated at

Žabout 10% of the observed geoid anomaly Dahlen,.1982 . If that estimate is approximately correct, the

interpretation of geoid anomalies can be simplified:the density contrast between upwelling and down-flowing mantle will be dominant and the strongpositive geoid anomaly over the North AtlanticŽ .Haxby, 1985; Heirtzler and Frawley, 1994 can betaken as evidence for relatively dense mantle andsub-axial downflow.

1.1.11. Seismic indicators of sub-ridge conÕergentflow

Evidence favoring the proposed viscous boundarylayer model and sub-ridge convergent flow is pro-vided by recent summaries of upper mantle anisotro-py. Interpretations of Pacific seismic velocities arecomplicated by the apparent existence of two distinc-tive patterns of seismic anisotropy. A principal pat-tern is clearly related to the EPR. As shown below, itprovides a key line of evidence regarding the direc-tion of mantle flow. A second pattern is apparently

Žrelated to an ancestral Mid-Pacific Ridge Section.2.5 .

Ž .Montagner and Tanimoto 1990 showed thatLove-waves and Rayleigh-waves exhibit maximalanisotropy below mid-ocean ridges, notably belowthe EPR and below the ridge segment south ofAustralia, and that their fast propagation directionsare nearly ridge-normal. They noted, as a puzzling

w xresult op. cit., p. 4816 , the similar fast directions ofLove-waves and Rayleigh-waves. Equally puzzling,and an important clue to the sub-EPR flow regime, is

that the Pacific Love-waves have fast propagationŽdirections parallel to those of P-waves cf. Hess,.1964; Raitt et al., 1971; Chesnokov, 1973 , whereas

they should be 908 apart if within the sameŽanisotropic medium Montagner and Nataf, 1986, p.

.517 . In a zone of shear flow, the fast direction ofP-wave propagation is parallel to crystallographic aof olivine but the fast Love wave will be that polar-ized parallel to olivine a axis and propagating nor-

Ž .mal to a Fig. 3 .The proposed viscous boundary layer model of

the upper mantle provides a way of accounting forthe parallel P- and Love-wave fast directions below

Ž .the mid-ocean ridge. It is proposed Fig. 3b that Pn

paths sample uppermost sub-moho mantle, a LVZdominated by ridge-normal shear flow. The olivine a

Ž .axis is parallel to shear flow hex-a model , andconsequently the P-wave fast axis is orthogonal tothe ridge. Love-waves are proposed to traverse asomewhat deeper level, where the near-axial mantleis dominated by ridge-normal compression. In thatregime, the olivine b axis is parallel to direction of

Ž .shortening hex-b model . The a–c plane will beparallel to the ridge, and the polarization of Love-waves in that plane accounts for their fast propaga-tion direction, orthogonal to the ridge. The hex-borientation in a regime of uniaxial compression, witha and c randomly oriented in the plane normal to bŽ .Karato, 1987 , was confirmed by the experiments of

Ž .Kern 1993 , who reported the absence of splitting inS-waves traversing a mantle xenolith parallel to the

Ž .b axis 010 .Ž .In the proposed model Fig. 3 , the zone of

ridge-normal compression is a feature of near-axialupper mantle and there will be a change, with prox-imity to the axis, from shear flow to ridge-normalcompression. That change serves to account for anobserved 908 rotation, at many ridge sites, in theorientation of the fast wave of an SKS split-wavepair, from ridge-normal fast for long paths to ridge-

Žparallel fast for near-axial paths Hung and Forsyth,.1997 .

1.1.12. Sub-axial seismic ÕelocitiesA key test of upwelling versus downflow is pro-

vided by data regarding the velocity of teleseismicP-waves that follow steep, near-vertical paths in the

Žnear-axial region. At an upwelling flow corner up-


.welling leading to divergent lateral flow , the axis ofshortening will be vertical and the olivine b axisŽ .slowest P-wave velocity will have a vertical orien-tation, thus further decreasing the slow V of aP

vertical path within a hot rising plume. At a zone of

Žhorizontal convergent flow, on the other hand con-.vergence leading to downflow , the olivine b axis

will have a horizontal orientation at the flow cornerŽ .zone of uniaxial compression . The a and c axesŽ .fast and intermediate V , will tend to be randomlyP


oriented in a vertical plane, thus adding to the highP-velocity characteristic of a vertical path in colddownflowing mantle.

Ž .A critical P-velocity test Blackman et al., 1993was based on data from an array of five submersibleseismometers deployed across the MAR at 25-kmspacing. Blackman et al. noted that published mid-

Ž .ocean ridge models e.g. Phipps-Morgan et al., 1992led them to expect a significant decrease in seismicvelocities below the near-axial region, due to theeffects of higher temperature and partial melting in apresumed upwelling mantle plume. They found, con-trary to expectations, a broad region of early P-wavearrivals, up to 15% fast below the near-axial region,tapering to less than 1% fast between 30 and 60 kmoff axis. The investigators apparently did not realizethe fundamental significance of their discovery. Theyattempted to interpret the results in accord with theplate tectonics model by suggesting that the fastP-wave velocities of steep near-axial paths may re-flect anisotropy due to vertical orientation of thinfilms of a melt fraction in a rising plume. A follow-upstudy dismissed that possibility, however, and led theauthor to attribute the high P-velocities to extremeolivine-alignment anisotropy in the AupwellingB re-

Ž .gion Kendall, 1994 . The calculated 15% anisotropyis very high, however, as compared with the typical7% to 8% maximal anisotropy related to horizontal

Žmantle flow beneath oceanic crust Shimamura et al.,.1983; Kawasaki, 1986 .

A similar P-velocity experiment was recently re-Ž .ported for the EPR Toomey et al., 1998 . P- and

S-waves travelling in sub-horizontal, cross-axialpaths show increasing delays as the ridge axis is

approached, a result consistent with a shallow sub-ridge low-velocity zone. Results for steep P-wavepaths are similar to those of the MAR experiment;P-waves with near-vertical paths are relatively fast,progressively faster for the steepest paths. The au-thors concluded that the early arrivals are the resultof vertical alignment of olivine a axes in an up-welling zone. That conclusion is dependent on theassumption that the alignment effect dominates con-trary effects from higher temperature and partialmelting, typical features of an assumed upwellingplume. No such balancing of opposed density effectsis required for the proposed recycling model of

Ž .mid-ocean ridge basalt MORB volcanism. The ob-served fast velocities of steep-path P-waves belowthe ridge axis can be attributed to the additive effectsof olivine alignment and relatively cold downflowingmantle, together with possible effects from the ani-

Žsotropy of mylonites Cannat et al., 1991; Mc-.Donough and Fountain, 1993; Kern, 1993 , a pre-

dictable feature of downflow traction.Dimensions of the deduced broad zone of high

seismic velocity and deduced downflow can be esti-Žmated roughly from the observation Blackman et

.al., 1993 that the high velocities of steep-path P-waves taper off with distance from the axis butremain above oceanic average out to a distance of30–60 km. A similar width can be deduced from the

Ž .results of Bottinga and Steinmetz 1979 . Their P-wave data for off-axis moho-depth earthquakesshowed a general downward refraction toward theaxis for distances out to 30 km from the axial valley.

Ž .The MAR seismic array of Blackman et al. 1993did not include a seismometer site within 25 km of

Ž .Fig. 3. A Upper mantle seismic anisotropy from regionalization of Love-wave velocities to show the relationship of fast propagationŽ .directions to the mid-ocean ridge system heavy line , including East Pacific Rise and Southern and Indian Oceans. Redrawn from Fig. 5 of

Ž .Montagner and Tanimoto 1990 . Individual lines represent fast propagation directions and line lengths are proportional to the amplitude ofŽ . Ž .Love-wave anisotropy. B Block diagram viscous boundary layer model through near-axial mantle below a mid-ocean ridge to account for

the parallel P- and Love-wave fast directions in terms of olivine orientations at two levels, an uppermost sub-moho layer dominated by shearflow and an underlying layer dominated by massive body flow and ridge-normal compression. Olivine orientations are represented byquasi-hexagonal models, the Ahex-aB model for orientation within the shear zone, the Ahex-bB for orientation under uniaxial compressionm.

Ž .The Ahex-aB model, with crystallographic a as symmetry axis, is widely used in place of the true orthorhombic model upper block inset ,Žas a means of simplifying the interpretation of anisotropy Ave Lallemant and Carter, 1970; Nicolas et al., 1973; Nicholas and Christensen,´

. Ž1987; Christensen, 1984; Babuska and Cara, 1991 . Within the zone of shear flow the fast P-wave is ridge-normal, parallel to a Raitt et al.,. w x1971 . Within the zone of uniaxial compression the Ahex-bB model is appropriate. Crystallographic b axis 010 has a preferred orientation

Žparallel to the direction of shortening, and the a and c axes are randomly oriented in the plane normal to b Karato, 1987, p. 456; Ribe and. Ž .Yu, 1991; Borman et al., 1993; Kern, 1993, p. 252 . The fast Love wave vibrating in the a–c plane has a ridge-normal propagation

direction.


the median valley, and therefore no P-velocity datawere obtained for the postulated narrow tongue of

Ž .sub-valley low-velocity material Fig. 4 . A broadbut shallow low-velocity zone is a feature of alloceanic ridges and there are indications of a near-vertical low-velocity tongue below the ridge axis.

Ž .Toomey et al. 1998 reported low P-wave velocitieswithin a narrow axial zone less than 10 km wide, and

Ž .Evans et al. 1998 reported a narrow electricalconductor beneath the crest of the EPR. Wolfe et al.Ž .1997 used travel times of body waves from tele-seismic earthquakes to outline a narrow cylindricalzone of low P- and S-wave velocities, of 150–200km radius, extending to at least 400 km belowIceland. They attributed the low-velocity anomaly,2% for P-waves and 4% for S-waves, to a hot narrowplume of upwelling mantle but the plume model isopen to question on the basis of transition zone

Ž .deflections see below . It would seem appropriate toconsider an alternative, consistent with the proposedrecycling model of volcanism, that the low-velocitysub-axial anomalies, exemplified by that below Ice-

land, are due to serpentine rather than to elevatedtemperature and partial melting. The experiments of

Ž .Christensen 1966, Fig. 5 show that 10–11% ofserpentine in mantle peridotite at 10 kb will accountfor a 4% S-wave retardation, as observed belowcentral Iceland

1.1.13. Transition zone deflectionsSeveral investigators have re-examined the plume

hypothesis by seismic determination of depths to the410 and 660 km discontinuities. Phase equilibrium

Žinvestigations Navrotsky, 1993; Bina and Helffrich,.1994 attribute the discontinuities to phase transi-

tions that have opposite Clapeyron slopes, such thathotter than average transition zone temperatures areindicated by downward deflection of the 410 andupward deflection of the 660 discontinuity, i.e. bytransition zone thinning. This can be referred to asthe Aopposed deflection modelB for the A410B andA660B discontinuities. Opposite deflections in thereverse sense, and transition zone thickening, arepredicted for regions of cold downflow.

Fig. 4. Hypothetical model of proposed mid-ocean ridge dynamics and structure, including sub-ridge convergence and downflow. CrustalŽ .thickness is greatly exaggerated in order to show the structure: a AflexloadB syncline cf. Figs. 5 and 6 of former wide extent that reflects a

Ž .crestward increase in volcanic loading and subsidence. Mantle boundary layer and isotherm shapes are adapted unscaled from theŽ .sub-mountain downflow diagram of Houseman et al. 1981 , for the dynamics of a continental mountain range, assumed to have a structure

Ž .similar to the oceanic range. The LVZ mixing zone is assumed to be equivalent to the seismic low velocity zone see text . The boundarybetween mantle and sub-axial wedge of subducted crust is a proposed zone of dehydration melting. Crustal layerings fine dashed lines;

Ž . Ž .isotherms non-specific smedium weight solid lines; Moho and metamorphic isograd sM; LVZ zone of crustrmantle mixings fine dotpattern; sub-axial zone of dehydration meltingsmixed dot pattern; basalt depleted residuumsXX pattern.


A recent summary of worldwide data for PdsŽphases P to S conversions at depth d below the. Ž .receiver Chevrot et al., 1999 showed very small

differences in the transition zone deflections, pre-sumably reflecting very slow mantle flow. Theynoted a deficiency in oceanic data, and it seemsdesirable to re-examine the evidence for oceanicregions in which strong downflow is indicated bystrong positive geoid anomalies. The Pacific trench–arc systems show downward deflection of the A660BŽCastle and Creager, 1998; Flanagan and Shearer,

.1998 , consistent with the Aopposed deflectionBmodel. Results are mixed, however, for the axialzone of the mid-ocean ridge system below the EPR,which lacks a strong geoid anomaly. No transitionzone deflections were found by two independent

Žinvestigations Lee and Grand, 1996; Shen et al.,.1998b . Below Reykjanes Ridge and Iceland, marked

by a strong positive geoid anomaly, a similar seismicstudy showed transition zone thinning and led the

Žauthors to deduce a deep-source mantle plume Shen.et al., 1998a . That interpretation is puzzling, how-

ever, because the departures of the sub-Icelandic 410and 660 transitions from a standard Earth model arenot anticorrelated, as expected for the Aopposed de-flectionB model. The P410s waves are delayed 7 s,on average, and the P660s waves are delayed an

Ž .average of 4.7 s Shen et al., 1996 . The indicateddownward deflection of both transitions is consistentwith relatively hot mantle at the 410 level andrelatively cool mantle at the 660 level. We are leftwith considerable uncertainty, but note that the pro-posed recycling model of oceanic volcanism offers asimple alternative: that the sub-Icelandic low veloc-ity mass is not a hot plume, it is a segment of theMAR in which the characteristic crustal recycling isaccentuated, with the result that there a zone ofcrust:mantle mixing is at the margins of a central

Ž .wedge of subducted crust Fig. 4 . In that model theelevated temperatures at the 410 level can be at-tributed to the exothermic hydration of olivine toserpentineq talc, while the 660 level, below theserpentine-forming zone, is dominated by colddownflowing mantle.

1.1.14. Mantle diapirs and uplifted old crustDiapirs of serpentinite and serpentinized mantle

peridotite are reported on the margins of the median

Ž .valley Cann et al., 1992 , near the ridge axisŽ .Bonatti, 1976 and along the margins of majorfracture zones, notably at the inside corners of RTI.Their tectonic association is consistent with the pro-posed recycling of hydrous crust into the uppermantle. Some of the transform-related diapirs formelongate transverse ridges, others generate separateuplifted massifs, typically 10 to 20 km in plandimensions. In one form or the other, they are acharacteristic feature of major Atlantic fracture zones:

ŽVema FZ Bonatti and Honnorez, 1971; Auzende et. Ž .al., 1989 , St. Paul’s FZ Melson et al., 1972 , Kane

Ž . ŽFZ Auzende et al., 1994 , Atlantis FZ Blackman et. Žal., 1998 , and Romanche FZ Bonatti et al., 1976,

.1979 . Generation of the mantle diapirs can be at-tributed to local recycling of hydrous crustal materi-als and consequent serpentinization of mantle peri-dotite. Local recycling of surface materials, in accordwith the proposed model, is consistent with theoccurrence, on Vema transverse ridge, of ultrama-fic-carbonate breccias in which the carbonate cementhas the carbon, oxygen and strontium isotope ratiosof present-day oceanic carbonates, and some of the

Ž .samples contain tests of forams Bonatti et al., 1974 .The size and shape of the mantle diapirs and the

ages of their cap rocks provide information regardingthe mechanism and time of their formation. Those inthe North Atlantic, relatively protected from theMesozoic surge of mantle flow, have nearly equant

Žplan dimensions, typically less than about 15 km see.examples in Fig. 2 . In the South Atlantic and Indian

Oceans, judged to be more affected by the Mesozoicsurge, some of the diapirs are elongate ridges, ex-tending for hundreds of kilometers along the bordersof major fracture zones, traces of the margins offormerly active convective rolls. It is implied thatthere was active recycling of hydrous crust alonghundreds of kilometers of convective rolls in thesub-oceanic mantle and that the lengths of the di-apiric ridges can be taken as a minimum ocean widthat the time of roll margin subduction and diapirformation. The classic Vema diapiric uplift is a370-km-long transverse ridge that rises 2 to 4 kmabove the Vema FZ valley in the equatorial Atlantic.Where a nearly complete crustal section is exposed,for example on the steep, transform-facing walls of

Žthe transverse ridge, submersible traverse of.Auzende et al., 1989 , the base of the uplift is


Žcomposed of serpentinized mantle peridotite up to 1.km exposed . The peridotite base is overlain by a

Ž .relatively thin crustal section cf. Fox et al., 1973 ,comprised of about 500 m of gabbro cut by ridge-parallel dolerite dykes that merge into 700–1100 mof a dolerite dyke complex, capped by up to 800 mof eruptive basalt, and 500 m of shallow-water lime-

Ž .stone Bonatti and Honnorez, 1971; Ligi et al., 1994 .The highest parts of the Vema transverse ridge, nowat 600–800 m depth, show evidence of former emer-

Ž .gence or shallow water conditions Table 1 , includ-ing a flattened, wave-cut base below a limestone cap.Bonatti and Honnorez dredged rounded pebbles ofhighly weathered basalt from the Vema ridge crestand noted additional evidence of subaerial exposure,based on dissolution features and local recrystalliza-tion of the limestone to a low-Mg calcite with low

18 Ž .d O y3.8‰ , characteristics of recrystallization bymeteoric waters. The age of the Vema limestone cap

Ž .apparently is open to question. Honnorez et al. 1975had the enclosed fossils examined by several paleon-tologists and tentatively assigned a Mesozoic age,

Ž .but a later paper Bonatti et al., 1994 favored aTertiary age.

The Romanche transverse ridge is an extensivediapiric ridge along the north side of the RomancheFZ, which offsets the MAR axis by about 900 km.About 100 km east of the eastern ridge–transform

Ž .intersection RTI , several of the Romanche peaks

are capped by sedimentary sequences, delineated byseismic reflection profiles. Dredging on the most

Ž .easterly peak ca. 158W yielded ventifact basaltpebbles and samples of a sedimentary sequence thatincludes Lower Cretaceous pelagic limestones over-lain by Paleocene to Eocene terrigenous siltstonesŽ .Bonatti et al., 1996 . Farther to the west, closer tothe eastern RTI, several peaks of the Romancheridge are flattened by wave erosion and a peak at17–17.88 W is capped by a shallow-water carbonate

Žplatform and reef deposits of Miocene age Gasperini.et al., 1997 . In view of the discovery of Cretaceous

limestones near the eastern end of Romanche ridge,it would seem important to confirm the age of theVema limestone by drilling to get fresh samples fordetailed paleontology and stratigraphy. It is sug-gested that ages of the sedimentary caps can bebracketed by their strontium isotope ratios, poten-tially diagnostic because of the major post-Jurassicincrease in the oceanic 87Srr86Sr ratio, from a mini-mum of 0.7068 to a present day value of 0.7092Ž .Veizer, 1989; cf. Wickman, 1948 . Further indica-tions of ancient rocks on the Mid-Atlantic ridge, andof non-spreading of the oceanic crust, are providedby a Jurassic 40Ar–39Ar age for a basalt dredged from

Žthe median valley–Atlantis FZ junction Ozima et.al., 1976 , by a K–Ar age of 835 Ma on hornblende

mylonite from St. Paul’s rocks, small islets repre-senting the summit of the St.Paul’s tranverse ridge,

Table 1Former shallow water or emergent sites, MAR

Ž .Site Location Depth m Indicators References

Ž .DSDP 114 59.98N 1927 H 7 v. 12, p. 313, 320, 326Ž .DSDP 384 40.48N 3903 S, V, W, H 20 Shipboard Scientific Party, 1976, v. 43, p. 107, 131, 776

DSDP 407 63.98N 2472 P, W Shipboard Scientific Party, 1978b, v. 49, p. 21, 39, 45, 51aDSDP 409 62.68N 832 V Shipboard Scientific Party, 1978b, v. 49, p. 161– 168, 174, 186

DSDP 556 38.98N 3672 P, W v. 82, p. 61, 70, 72DSDP 558 37.88N 3754 P, W v. 82, p. 127, 143

Ž .208N 1920 W Auffret et al., 1991Ž .26–278N )3000 ES Tucholke et al., 1997Ž .Romanche 0.48N, 178W 800 P, TS Bonatti and Chermak, 1981; Bonatti et al., 1992Ž .E. Romanche 0.48N, 158W 2270 E, CS Bonatti et al., 1996; Gasperini et al., 1997Ž .Vema FZ 118N, 42–458W 600 E, P, TS, W Bonatti et al., 1971, 1994; Honnorez et al., 1975

Ž .Es ridge top flattened by wave erosion; Hsmajor hiatus in basal sedimentary section missing time span, Ma ; Psbasaltic pebblesŽ . Ž .ventifact on Romanche ; Vshighly vesicular basalt; CSsCretaceous shallow water sediment; ESscanyons, trellis drainage sub-aerial? ;Wsweathered and oxidized basalt; TSsTertiary shallow water sediment.

a Pipe vesicles at site 409.


and by the discovery of Paleozoic and older zirconsŽ .U–Pb ages 330 and 1600 Ma in uplifted lower

Ž .crustal gabbros near the Kane FZ Pilot et al., 1998 .Some of the mantle diapirs have been completely

unroofed, for example along the Kane and Atlantisfracture zones, with the result that the structure ofthe near-transform mantle is exposed. Notable detailis presented in a recent acoustic map of the Atlantis

Ž .FZ and its RTI Blackman et al., 1998 . Exposedmantle rocks of unroofed diapirs adjacent to thetransform exhibit a pronounced transform-parallel

Ž .structure, presumably a shear zone Fig. 2 , notablebecause it is at 908 to the regional ridge-parallelstructure of the oceanic crust. Similar examples ofexposed transform shear zones, from the Kane FZŽ .Atlantic , and from the Indian Ocean, are shown in

Ž .a short article by Mitchell et al. 1998 . The detailedŽ .chart of the Atlantis FZ Blackman et al., 1998

provides indication of a probable mechanism of un-roofing of the diapiric domes. The uplifted mass atthe eastern ridge-transform intersection is borderedby an apparent crustal mass low on its eastern slope,in a position and altitude that lead one to suggestsliding of the crustal cap into the adjoining axialvalley. One possibility, consistent with identificationof the diapirs as serpentinized peridotite, is that therising diapir brought serpentine through its stabilitylimit, so that the pressure effect of released watervapor promoted sliding of the overlying crust. Dehy-dration of serpentine to olivineq talcqwater occursat an invariant point at 7008 and 21 kbar, but atlower temperatures for either higher or lower pres-

Ž .sures Ulmer and Tromsdorff, 1995 . It is accompa-nied by tectonic weakening and the development oflinear viscous rheology, largely accomplished byslippage along shear zones lined with ultrafine olivineŽ .Rutter and Brodie, 1988 .

The above examples of mantle diapirs are mainlyfrom the Atlantic ridge system. Similar diapiric ridgesof serpentinized peridotite occur in the Indian Ocean,

Žfor example along the Atlantis II FZ Shipboard.Scientific Party, 1989, p. 42 . The longest one, Nine-

tyeast Ridge, is analogous to the Vema and Ro-Žmanche transverse ridges in the Atlantic Sclater and

.Fisher, 1974 but is now displaced far to the North,well away from its original oceanic ridge locationŽ .Liu et al., 1983 . It was proposed in early reportsŽ .Bowin, 1973; Shipboard Scientific Party, 1974 to

be a gravitational uplift of serpentinized peridotitewith a transform origin, and that seems evident fromits 908 orientation relative to bordering crustal struc-ture, and from the observation that magnetic anomalyages increase in opposite directions on either side of

Ž .the ridge Souriau, 1981; Deplus et al., 1998 . North-ward migration of Ninetyeast Ridge is indicated bythe observation that nannofossils from Broken Ridge,

Ž .at the southern end 318S , are high latitude formssimilar to assemblages from the Falkland PlateauŽ .Weissel et al., 1991, p. 141 . Available detailsregarding the basaltic cap and sediment cover of

Ž .Ninetyeast Ridge papers in Weissel et al., 1991lead to the suggestion that it can be taken as a crustalsection across the ancestral Southeast Indian Ridge,from the intersection with the former axial zoneŽ .now represented by Broken Ridge, at 318S , to thenear-equatorial northern end. The uplifted base ofNinetyeast Ridge is deduced, from seismic and grav-ity surveys, to be composed of serpentinized peri-

Ž .dotite Bowin, 1973; Souriau, 1981 , and the caprocks are MORB and overlying sediments that rangefrom shallow-water Eocene beds near the southernend to Cretaceous chalks near the northern limit.Radiometric 40Arr39Ar ages of uppermost basalt showa range of 42 Ma at the southern, near-axial end, to

Ž .88 Ma in the North Duncan et al., 1991 . CrustalŽ .rocks at a southerly location ODP Site 756 and a

Ž .central location ODP site 757 are oxidized sub-aerial basalt, indicating former exposure of much ofthe Ninetyeast Ridge.

1.1.15. Clues to sub-fracture zone and sub-axialrecycling

Sub-fracture zone and sub-axial recycling of hy-drous materials is evidenced by the common occur-

Ž .rence of serpentinized mantle diapirs see above andby the associated emissions of hydrogen and methane,by-products of the exothermic serpentinization of

Ž .mantle peridotite Charlou et al., 1997 . The exother-mic character of the process is consistent with theobserved hydrothermal fields above outcrops of ser-

Ž .pentinized mantle peridotite. Douville et al. 1997described an example, a hydrothermal field in anultramafic environment at a fracture zone–axial zoneintersection. There is no indication of recent volcan-ism at that site and the authors suggested that the

Ž .hydrothermal circulation at exit temperature 3608C


is activated by the exothermic serpentine-formingreaction.

The sub-ridge LVZ, here proposed to result fromcrustal recycling at roll margins and the ridge axis, isseveral hundred kilometers wide below the EPR andexhibits maximal Rayleigh-wave retardation at 15–70km depth, which the authors attributed largely to

Ž .partial melting MELT Seismic Team, 1998 . Asimilar LVZ is a typical feature of other ridges, and adeeper LVZ mass is reported below central IcelandŽ .Wolfe et al., 1997 . The authors attributed the Ice-landic anomaly to a hot mantle plume but the ob-

Ž .served sinking of Iceland Section 1.2.3 leads to thepossibility that the LVZ mass is formed by recyclingof hydrous crust and resultant generation of serpen-tinized mantle peridotite. Local recycling of old crust

Ž .and intercalated sediments Section 2.3 is evidencedby an array of isotopic fingerprints.

1.1.16. Vesicle pressure in MORB glassŽ .Moore et al. 1977 measured volume expansion

of carbon dioxide released from pierced vesicles inMORB glass slabs immersed in glycerine and founda simple positive correlation between vesicle pres-sure and collection depth over a depth range 715–

Ž4828 m volume expansion estimated at 20"5 times.the collection depth, in km . If the represented basalt

had erupted at the axial zone and reached presentŽ .depths by ocean floor spreading PT model , the

internal vesicle pressure might be expected to varywithin narrow limits, due to a range of shallowdepths at the postulated axial source, but there wouldbe no reason to expect vesicle pressure to show asystematic relationship to collection depth. The re-sults are consistent with the proposed widening ofthe zone of active volcanism, a regime in whichvolcanism presumably occurred over a wide range of

Ž .ocean depths Sections 3.1.3 and 3.2.1 . It would beof interest to apply the vesicle pressure measure-ments to the popping rocks collected from 3400 to

Ž3600 m depths on the MAR Sarda and Graham,.1990; Pineau and Javoy, 1994 . Their high vesicle

pressures are consistent with deep, off-axis extru-sion.

Some of the above indicators of the direction ofmantle flow can be accounted for in several waysand can, with difficulty, be adapted to the platetectonics model. I propose that, on balance, they

should lead to its rejection and tend to favor theproposed alternative model. It is suggested that read-ers may wish to use their own interpretations of thevarious clues and summarize conclusions in the tablein Section 1.3.8.

1.2. Oceanic mantle dynamics and surface expres-sions

1.2.1. A Õiscous boundary-layer model of upper man-tle

In the proposed model the upper mantle functionsas a viscous boundary layer that merges gradually

Žfrom sub-lid shear flow to deeper body flow Ribe,.1989; Keith, 1993 . The term Asub-lidB is used in

preference to sub-crustal, in order to allow for thepossibility, brought out by the occurrence of my-

Žlonite structure in lower crustal gabbro Cannat et.al., 1991 , that the lower crust may be a part of the

boundary layer. The seismic evidence is generallyconsistent with the proposed viscous boundary layermodel. Seismic anisotropy in oceanic upper mantle isgenerally attributed to olivine orientation by shear

Žflow in uppermost sub-moho mantle Hess, 1964;.Bottinga and Steinmetz, 1979; Ribe, 1989 .

Anisotropy deduced from P waves, constrained tonŽnear-surface paths to about 50 km beneath ridge

.flanks , typically shows the V -fast direction parallelP

to fracture zones and transforms, consistent with amantle flow direction that is ridge-normal or highly

Žoblique to the ridge axis Morris et al., 1969; Nicolasand Christensen, 1987; Nishimura and Forsyth,

.1989 . S-wave tomography at greater depths, begin-ning between 200 and 370 km depth, shows manyexamples of a 908 change in azimuth, from ridge-normal to ridge-parallel fast-direction of polarizationŽ .Montagner and Tanimoto, 1990 . Similar ridge-parallel fast directions at depth were reported in anearlier study of SS-S differential travel times in the

Ž .North Atlantic Kuo et al., 1987 . The authors at-tributed that orientation to North–South ridge-paral-lel mantle flow. Alternatively, an abrupt downwardchange from ridge-normal to ridge-parallel polariza-tion of the fast ray can be accounted for by theproposed viscous boundary layer model of conver-

Ž .gent flow Fig. 3 in which olivine orientation inuppermost mantle reflects preferred alignment ofcrystallographic a parallel to the direction of shear

Ž .flow hex-a model but at deeper near-axial levels is


due to alignment of crystallographic b parallel to theŽ .direction of shortening hex-b model , within a zone

of convergent flow and uniaxial compression.

1.2.2. ConÕectiÕe driÕeMantle convection, primarily due to heating from

within and below, is subject to certain focusingeffects due to characteristics of the upper and lowerboundary layers. The results of experimental convec-tion, together with the above-noted clues to thedirection of mantle flow, indicate that upwellingtends to be focused beneath thermal barriers, mainlybeneath continental cratons, and that a principal re-sult is circum-continent outflow and a sub-oceanicsystem dominated by sub-ridge convergent mantle

Ž .flow and sub-axial downflow Fig. 4 . Interpretationsof the surface expressions of mantle flow are con-strained by the simple but fundamental experiments

Ž .of Griggs 1939 . He carried out a series of viscousflow tests with a deformable viscoelastic crust sup-ported on a moving fluid substrate analogous toflowing mantle and showed that the typical surfaceexpression of upwelling and divergent flow is abasin or depression, a reflection of stretching and

Žthinning of the deformable surface layer cf. geoid.anomalies, Heirtzler and Frawley, 1994 . Con-

versely, the surface expression of a zone of conver-gent and downwelling flow was shown to be athickened and uplifted surface layer, analogous to aplateau or mountain range. The Griggs model can beapplied to interpretation of ocean ridge dynamics butis oversimplified in that it does not allow for iso-static uplift due to formation of a low-density crust–mantle mixture at sites of cell boundary downflow.That effect is positive, however, in the same sense asthat due to crustal thickening, and is presumed to beparticularly significant in oceanic regions because ofthe subductibility of metamorphosed oceanic crustŽ .see Section 2.1 . Gabbro of the lower oceanic crustpresumably is altered to amphibolite, enters the gar-net stability field at upper mantle P–T conditionsŽPoli, 1993; Wyllie and Wolf, 1993; Bohlen et al.,

.1994 and overlaps the density of serpentinized up-per mantle peridotite.

The boundaries of convective rolls constitute adownflow regime of secondary convection. In themid-ocean ridge setting they are represented by frac-ture zones and transforms and the roll boundary

recycling of hydrous crustral materials is proposed toyield serpentinization of mantle peridotite and het-erogeneous crust–mantle mixtures, identified as thesub-ridge LVZ. That model of LVZ formation leadsto the generalization that the lateral extent and depth

Ž .of the LVZ Zhang et al., 1994 is related to theextent and dimensions of the convective rolls.

The sub-axial zone is proposed to be a zone ofrelatively deep recycling where the hydrous LVZmixture is subject to volatile-promoted partial melt-

Žing, with resultant generation of basalt Section.1.1.15 . A postulated side effect is the development

Ž .of an underlying sub-axial mass of residuum Fig. 4that is depleted in basalt-forming constituents and is

Ž .also relatively viscous depleted in volatiles andŽ .relatively cold depleted in radioactive elements .

Accumulation of a cold, viscous residue during pro-longed steady-state convection is proposed to focuscontinued, relatively strong downflow at the axis.

Where a cold residue is well developed, the axialzone is characterized by a strong positive geoidanomaly and by well-developed median valleys, forexample on the northern MAR, they are taken asindicators of strongly focused downflow. On theother hand, a relatively young sector of the ridge,such as the displaced EPR, will not have a welldeveloped cold residue at depth below the axis, andthe dynamic system will be dominated by convergentflow rather than by focused sub-axial downflow.That difference may account for the lack of medianvalleys and of a strong positive geoid anomaly onthe EPR.

1.2.3. Icelandic structureA principal source of evidence regarding oceanic

crustal structure and mid-ocean ridge dynamics is theexposed Icelandic sector of the MAR. Many of thosewho favor the plate tectonics model regard Icelandas a mantle plume site that cannot be taken torepresent the typical axial zone of the mid-oceanridges. The local plume model is open to question,however, on the grounds that Icelandic heat flow is

Žnot above the ridge average if active volcanic areas.are omitted . Furthermore, Iceland is the culmination

of a gradual northward shallowing of the Ridge inthe region to the north of the Charlie–Gibbs FZ,over a total distance of about 1500 km. It is clearlyunreasonable to dismiss the evidence at the Icelandic


culmination of that long rise, the only place wherewe can see the structure.

The major structure of the near-axial zone, asŽ .exposed in Iceland Fig. 5 , is that of a large vol-

canic syncline characterized by volcanic formationsthat dip inward toward the axial zone and systemati-

Žcally thicken in the downdip direction Walker, 1960;.McDougall et al., 1984 . Tertiary basalts on either

side of the active volcanic zone commonly dip 20–Ž308 below the edge of the younger flows Saemunds-

.son, in Vogt and Tucholke, 1986 . There are someirregular structures on the inner side of the Icelandic

Ž .regional bend FT of Fig. 5 , and some anticlinalflexures bordering the axial volcanic zoneŽ .Saemundsson, in Vogt and Tucholke, 1986 , butthey are minor structures relative to the maingeosyncline. Over most of Iceland and over a totalexposed width of about 500 km, the dips typicallysteepen toward the axial zone and also steepen grad-ually with depth. The observed synclinal structure

Ž .Fig. 5. Structural outline map of Iceland, redrawn from Saemundsson 1986 with added indicators of the near-surface measured stressorientation. FTsRegion of irregular folding of Tertiary volcanic series; TsThingvellir, classic rift valley site on the landward extensionof the Mid-Atlantic Ridge; A and KsAskja and Krafla, recently active volcanoes in the northern volcanic zone; Bssite of Borgarfjordur

Ž .test hole see text ; Open arrows show direction of maximum horizontal compression at the following stress measurement sites, whereŽ . Ž . Ž . Ž . ŽsH )1.5 sH : Hvalnes H data from Hast, reported in Haimson and Voight, 1977 ; Reykjavik R Haimson and Voight, 1977;max min. Ž . Ž . ŽSchafer, 1979 ; Storu-Tjarnir ST Haimson, 1979 ; S7, S13. Unnamed sites of measurements by Schafer 1979 and personal communica-

.tion .


was incorporated in an early schematic modelŽ .Bodvarsson and Walker, 1964 and in later quantita-

Ž .tive models Palmason, 1986 that postulated spread-ing combined with a systematic riftward increase inthe balanced rates of volcanism and subsidence.Abundant evidence for axial zone sinking comesfrom well-exposed cliff sections, some over 1 km

Ž .high Saemundsson, 1986, p. 69 , and from drill holelogs and geodetic data, taken to indicate near-axial

Žsubsidence of about 6.5 mmryear Sigmundsson and.Vadoh, 1996 . Seismic reflection profiles across the

axial zone in SW Iceland show steeply dippingŽlayers that can be traced to 15 km depth Zverev et

.al., 1980; Palmason, 1986 . Recent seismic investiga-tions show that the Icelandic crust reaches maximumthickness of 37–44 km in central Iceland, and isthinner below the neovolcanic zones on either sideŽ .Allen et al., 1999; Derbyshire et al., 1999 . Theauthors proposed a non-specific relationship betweenthick crust and a hypothetical mantle plume, but theIcelandic structure makes it evident that thick crust isa feature of the center of the Icelandic geosyncline,and represents a balance between sinking and re-

Ž .gional volcanic loading Section 1.2.4 .

1.2.4. Tilted crust and a A flexloadB modelIn the proposed mid-ocean ridge model, the struc-

ture of the crestal zone is attributed mainly to gravi-tational deformation, a result of two types of loading:

crestward-increasing volcanic loading and down-ward-increasing metamorphic densification. The re-sultant volcanic syncline is designated a AflexloadBstructure in order to emphasize the effects ofgravity-driven flexure on the crustal stress regimeŽ .Fig. 6 . There are indications that major flexloadsubsidence is not restricted to the near-axial zone butis broadly effective beneath the ridge flanks andformerly extended over the full width of Mesozoic toEarly Tertiary ocean ridges. Multichannel seismicreflection surveys of the central North Atlantic showmid-crustal reflectors that dip systematically toward

Ž .the axis Morris et al., 1993 . The character of thedipping structures and their similarity to Icelandicstructure is most clearly evident on ridge-normalprofiles that are remote from the structural complica-tions adjacent to major fracture zones. It seemslikely, from the available seismic reflection profilesŽ .Morris et al., 1993, fig. 6B that the near-fracturezone complications are due to a fracture zone-centered subsidence regime, similar to that beneaththe ridge, a regime that would be expected from theproposed generation of fracture zones at the bound-aries of convective rolls. Well defined examples ofridgeward-dipping structures, free of fracture zonecomplication, are provided by profile 711 of Morris

Ž . Ž .et al. 1993 their figs. 3, 4, 10a and 12a and by theŽ .Fig. 3 profile of Mutter and Karson 1992 . Similar

dipping reflectors, with dips of 108 to 408 toward the

Ž .Fig. 6. Idealized section of the subsiding crestal syncline of a mid-ocean ridge in the plane of a Transform offsetsT to show the proposedAflexloadB structure: downward flexure of volcanic formations and the consequent change from horizontal compression at shallow depth to

Ž .horizontal extension in the zone below the level of neutral stress LNS , a depth zone that defines the upper limit of a sheeted dike complexŽ .see text . The double line represents a zone of Astress transitionB that includes LNS and overlying crossovers of vertical stress withmaximal and minimal horizontal stress. LNB is the level of neutral buoyancy, a depth of balance between the density of basaltic magma andthe mean bulk density of upper oceanic crust.


ridge axis, are evident in seismic profiles in theŽ .central Pacific Eittreim et al., 1994 , and in Creta-

Žceous lower crust of the northwest Pacific Reston.and Ranero, 1999 . Mutter and Karson interpreted

the dipping reflectors of the MARK area of theNorth Atlantic as major detachment faults that pene-trate the entire crust. There is evidence, however,that the dipping reflectors are similar to the exposeddownwarped structures of Iceland, i.e. they are pri-mary crustal layers that have been tilted under vol-canic load. Indicators of widespread tectonic tilting

Žinclude skewed magnetic anomalies 33 and 34 Fig..8 and sediment interlayer dips, for example at near-

Ž .axial Atlantic sites 332, 395, 410 and 412 Fig. 8Žand at flank sites 417 and 418 Hall and Robinson,.1979; Donnelly et al., 1980, p. 75 .

Ž .Verosub and Moores 1981 suggested that someof the anomalously shallow magnetic inclinations inDSDP holes can be attributed to tectonic tilting in

Ž .the range 50–708. Cande and Kent 1985 noted thatmagnetic anomalies at the selected sites of Verosuband Moores are not skewed, as might be expected fortilted sections. Using the well-defined anomaly acrosswestern Atlantic Site 417 as an example, they showedthat the observed magnetic profile is closest to themodel with zero tectonic rotation. The deduced tilt-ing at Site 417 had been confirmed, however, by abedding plane dip of 558 recorded for interlayeredlimestone within basalt near the bottom of the holeŽ .Donnelly et al., 1980, p. 75 . Cande and KentŽ .1985, p. 4649 proposed that the apparent conflictcan be resolved by assuming that the oceanic mag-netic anomalies, mostly non-skewed, are not due tooriginal magnetization of basaltic upper crust, asformerly believed, but mainly to a deeper crustalsource that acquired remanent magnetism after tilt-ing, presumably during slow cooling through theblocking temperature.

Relatively rapid pre-metamorphic subsidence isevident in Iceland, judging from the observation thatsub-horizontal zeolite zones cut across the stratigra-

Žphy of the tilted volcanic pile Walker, 1960; Mc-Dougall et al., 1984; Saemundsson, in Vogt and

.Tucholke, 1986 . At deeper levels, sub-horizontalseismic reflectors have been attributed to highergrade metamorphic transitions within the crust, in-cluding the zeolite–greenschist transition and thelayer 2–3 boundary, tentatively equated with a con-

currence of gabbro intrusion and the amphiboliteŽ .isograd Magde et al., 1995 . The evidence leads one

to favor a process of progressive metamorphismsuperimposed on a volcanic pile subject to crestwardincrease in the rate of loading and subsidenceŽ .AflexloadB model, Fig. 6 . It seems likely that theseismically defined Moho in oceanic regions is a

Žcombined metamorphic and intrusive transition cf..Hacker, 1995 and that metamorphosed lower crustal

Žmaterials garnet amphibolite, pyroxenite and.metagabbro, for example eventually pass through

that transition and become incorporated in the lateralflow and downwelling of the upper mantle boundary

Ž .layer Keith, 1993 .The proposed mechanism is similar to the conti-

Ž .nental range model of Houseman et al. 1981 andŽfollows the suggestion of Platt and England 1994, p.

. Ž307 that Athe lower part of the lithosphere con-.tinental or oceanic is probably removed intermit-

tently by convection . . . B The proposed ocean ridgemodel differs from the above cited models in thateffective traction is suggested to involve hot creep inuppermost mantle and, to a lesser extent, in lowercrustal rocks. In contrast with continental crust, pre-sumably subductible only during strong surges ofmantle flow, relatively dense metamorphosed oceaniccrust can continue to subside during steady-stateconditions and will become an integral part of thebalanced process of subsidence, lateral transfer andrecycling.

1.2.5. Rock stress, geodetic and global positioning( )system GPS data and earthquake mechanisms

Evidence of axial zone dynamics is provided bystress measurements, by earthquake focal mecha-nisms and by GPS and geodetic measurements ofrelative crustal movements. Direct measurements of

Žcrustal stress in Iceland, some by overcoring Hast,.1973; Schafer, 1979 , some by hydrofracturing,

Ž .Haimson and Voight, 1977; Haimson, 1979 showedthat maximum horizontal compressive stress at shal-low depths on the borders of the active volcanic zoneis oriented approximately perpendicular to the axial

Ž . Ž .zone Fig. 5 cf. Palmason et al., 1979, p. 63 .Similar ridge-normal maximum stress has been mea-sured in DSDP boreholes 395A and 504B on the

Ž .MAR Moos and Zoback, 1990 . The significance ofthe direct stress measurements has subsequently been


dismissed or questioned, on the grounds that they arein conflict with the focal mechanisms of the greatmajority of axial zone earthquakes, which show thatminimum horizontal stress at earthquake depth is

Ž .oriented normal to the ridge axis Einarsson, 1991 .It seems likely, however, that both methods yieldmeaningful results and that the apparent conflictreflects differences of depth. In the proposed

Ž .AflexloadB model Fig. 6 , the stress measurementsŽrepresent a shallow compressional level see exam-

.ples in Fig. 5 , whereas the earthquake mechanismsthat indicate ridge-normal extension represent deeperlayers, below LNS, the level of neutral stress. In anelastic plate under flexure LNS, the plane betweenshortened and stretched portions is bent withoutchange of length. In competent adjacent layers, thecompressional and extensional forces increase as afunction of distance from LNS and are proportional

Žto the radius of curvature of the flexure Engelder,.1993 .

The depth of transition from compression to ex-tension, estimated from the upper limit of the seismiczone, is about 2 km in the Reykjanes Peninsula, a

Ždirect northward continuation of the MAR Klein et.al., 1977 appreciably shallower in the Tertiary lavas

of eastern Iceland, judging from stress measurementsŽin a research borehole, B of Fig. 5 Haimson and

.Rummel, 1982 . On the Mid-Atlantic and IndianOcean Ridges, typical axial earthquakes are normalfault events centrally located at 1 to 6 km depthsbelow the deep inner floor of the median valleyŽ .Huang and Solomon, 1988 .

Many recent investigations have used the GPS toevaluate intercontinental movements and to confirmthe ocean-floor spreading deduced from interpreta-tion of the oceanic magnetic anomaly pattern. Localtrans-rift GPS measurements, particularly in Iceland,are less ambiguous and more diagnostic because theyprovide a direct indication of spreading or non-spreading, whereas long-distance intercontinentalmeasurements fail to differentiate between plate mo-tions and the proposed alternative model: continentalmasses that are floating in the mantle and thereforesubject to differential motion as a result of changes

Ž .in the global convection pattern. Jonsson et al. 1995made repeated GPS measurement on a 59-km-longprofile across the volcanically active eastern rift inIceland, and demonstrated wide variability of trans-

rift motions, including significant contractions dur-ing the period 1967–1977, extensions during 1977–1994. A principal reason for that variability was

Ž .brought out by Pollitz and Sacks 1996 , who showedthat the principal motions are related to episodes ofvolcanic activity. They reported fissure opening, up-lift and local extensions up to 8 m in the Kraflasegment of the northern volcanic zone, during a1975–1984 episode of Krafla volcanic activity. Fol-lowing the period of active volcanism, the Kraflaarea was dominated by subsidence. Measurements

Ž .during 1992 to 1995 Sigmundsson et al., 1997demonstrated a 24 mmryear subsidence above themagma chamber, superimposed on 7 mmryear sub-sidence along the rift zone. During the 1975–1984time span of Krafla activity, the Askja volcaniccenter, 60 km to the south, was in a quiescent periodand was subject to large relative subsidence. It seemsevident that the deformations are local; they have thestop-and-go character of magma inflations and defla-tions and are not easily reconciled with the conceptof steady-state separation of thick lithospheric plates.

1.2.6. Axial and transform earthquakesA recent study of earthquake focal mechanisms

Ž .Julian et al., 1998 noted the occurrence of non-dou-ble-couple earthquakes associated with landslides andwith shallow zones of volcanic and hydrothermalactivity, including Krafla, on the Icelandic segment

wof the MAR. The authors suggested op. cit., 1998,xp. 565 that similar events are to be expected on

other segments of the mid-ocean ridge system. Thededuced non-double couple earthquakes do not pro-vide clear-cut evidence for or against the currentAspreadingB model, but they are consistent with theproposed collapse character of axial valleys.

A strong point of evidence in support of the platetectonics model is provided by interpretations of thefocal mechanisms of earthquakes that occur alongthe transform segments of fracture zones. The greatmajority of transform events yield a sense of shearmotion opposite to that expected for a simple offset

Žof the ridge crest Sykes, 1967; Einarsson in Vogt.and Tucholke, 1986 . An alternative hypothesis, con-

sistent with the proposed convergent mantle flow, isshown schematically in Fig. 7. Transform offsets ofthe ridge axis are a characteristic of regions wherethe trace of the ridge axis is curved, implying that


Fig. 7. Idealized relationship of fracture zones and transformsŽ .convective roll boundaries to the principal mid-ocean convection

Ž .cell boundary gray , and to offset axial valley segments in aregion where the mantle flow direction is non-orthogonal to theaxis of the mid-ocean ridge. The angular disconformity betweenaxial valley and transform boundary is exaggerated in order toemphasize the Atransform senseB of displacement of the axial

Ž .subduction zone S™S ™C toward the cell boundary in under-1

lying mantle. That displacement, increasing with depth, deservesconsideration as a possible mechanism of transform earthquakes.

the underlying mid-ocean convection cell boundaryis non-orthogonal to the general direction of mantle

Žflow, as indicated by fracture zones the boundaries.of convective rolls . At a transform, segments of the

Ž .axial valley S and near-axial crust are offset fromthe underlying curved cell boundary, and from thetypical rectilinear pattern, as a result of the difficultyof forcing the viscoelastic crust into a compoundcurve. As part of the geometric adjustment, the

Ž .sub-valley subduction zone S will be deflectedprogressively, with depth, from S toward S , for1

example, in a direction toward the actual sub-crustalcell boundary. That deflection has the relative mo-tion of the observed transform earthquakes. Relevantevidence consistent with that model includes indica-tions that the largest transform earthquakes typically

Žoccur near the center of the offset segment Burr and.Solomon, 1978 .

1.2.7. Flexural constraints on ÕolcanismIt is to be expected that the stress regime of the

proposed near-axial flexload syncline will constrainmany features of mid-ocean ridge volcanism, includ-ing dike and sill injection, and magma ponding incrustal reservoirs. The compressional and extensionalforces within a competent elastic layer under flexurewill affect s , the maximal horizontal stress. In ah

synclinal region of slight flexure, where verticalŽ .stress s is the largest stress throughout, dikes willv

tend to terminate upward by narrowing, beginning atLNS. On the other hand, in a syncline where flexuralstresses are significant and s is greater than s inh V

upper crust, the stress transition zone just above LNSwill be characterized by a crossover of horizontaland vertical stresses. Dikes will be favored belowLNS, and typically oriented parallel to the axis of the

Ž .flexload syncline cf. Fig. 5 , and there will be anupward, stress-controlled decrease in dike widths anddike abundance. That type of stress control is evidentin Iceland; the intensity of dike injection is greatest

Žat sea level and below 5% to 10% stretch at sea.level and diminishes upward to zero at the highest

Žexposures, at 1000 to 1500 m elevation Walker,.1960; Palmason et al., 1979 .

An additional effect of the flexload stress regimeis a tendency for sill injection and reservoir develop-ment at crossover depth, as rising magma encountersthe zone where s is less than horizontal stresses.v

Examples of the above injection styles are providedŽby detailed mapping in eastern Iceland Walker,

.1975 . Similar stress-controlled injection is to beexpected elsewhere on the mid-ocean ridge. Theimplied extensional stress regime at depths belowLNS is consistent with recent confirmation of asheeted dike complex in near-axial lower crust at

Ž .Pacific ODP Hole 504B Erzinger et al., 1995 . Itwill be evident that the flexload model serves toaccount for multiple dike injection and for the up-ward termination of dikes at a specific mid-crustaldepth. In contrast, the current, Apull-apartB model ofplate tectonics fails to account for either. After thefirst tensional failure, Apull-apartB stress would berelieved and there would be no pervasive tension, asrequired for a sheeted dike complex.

Ž .Prior investigators e.g. see Ryan, 1994 haveproposed magma ponding at a LNB, where the den-sity of basaltic magma matches the bulk density of


Ž .enclosing rock. Ryan 1994, p. 97 concluded thatthe zone of neutral buoyancy extends over a depthrange of 1–3 km, taken to include a deep LNB forrelatively dense parental magma and a shallower

ŽLNB for less dense tholeiitic magmas presumed.derivative . There are indications, however, that the

effective LNB is at shallower depths. Hooft andŽ .Detrick 1993 calculated the effective density:depth

curve for a segment of crust on the EPR, based onseismic velocity structure, together with in situ mea-surements of bulk porosity and density, and esti-mated that LNB lies only 100–400 m below the seafloor. Magma reservoirs at that depth will tend to beshort-lived, terminated by crystallization. They arethe probable cause of the elongated surface bulgesŽ .axial volcanic ridges that are observed along sub-ocean segments of the axial zone. Where MORBmagma reaches the surface, it typically is extrudedfrom a fissure on the crest of an elongate bulge. Itseems likely that fissure eruptions occur wheremagma pressure builds up in a shallow reservoir, sothat regional near-surface compression is counteredby bulging of the overlying crust. An additionalconstraint, related to the positive P–T slope of thebasalt liquidus, is that the magma must have appre-ciable superheat, above liquidus temperatures, other-wise the pressure increase required to bulge theoverlying crust will cause the magma to crystallize.

The proposed flexload model leads one to expectmore persistent magma reservoirs at LNS, the level

Ž .of neutral stress and stress transition Fig. 6 . Thattransition probably controls the depth of intrusive

Ž .sills in Iceland Walker, 1975 , and the 1–2 kmdepth of sill-like magma bodies imaged by multi-channel seismic reflection surveys across the EPRŽ . Ž .Wilcox et al., 1993 . Babcock et al. 1998 reporteda sub-axial magma lens or sill along the EPR at anaverage depth of 1335 m. There are indications thatthe stress transition is at a deeper level in IcelandŽ .Fig. 5 ; it is estimated at about 3 km, based on theindicated depths, from S-wave shadows, of the top of

Ža magma reservoir below Krafla caldera Einarsson,.1991 . The Krafla reservoir sits atop a thickened

lower crust containing a broad, high-velocity coneŽ .Menke et al., 1995 , possibly the accumulatedresidue of multiple fractionation episodes. Furtherevidence of the stress transition from tension at

Ž .depth to near-surface compression Fig. 6 is pro-

vided by observations during the 1975–1981 periodof volcanic activity at Krafla. Magma was dis-tributed, by lateral dike injection, northward andsouthward along the rift zone, for distances up to 60

Žkm from the central Krafla reservoir Einarsson and.Bransdottir, 1980; Sigurdsson, 1987 . Earthquakes

associated with that episode of lateral dike injectionŽwere in the depth range 3–4 km Einarsson, in Vogt

.and Tucholke, 1986 , consistent with the hypothesisthat upward dyke injection is constrained by thetransition from extension at depth to compression innear-surface layers.

On the eastern flank of the Icelandic volcanicsyncline, about 100 km to the east of Askja volcanoin the northern volcanic zone, stress measurements in

Ž . Žthe Bordarfjordur borehole B of Fig. 5 Haimson.and Rummel, 1982 indicate that the stress crossover

Ž .is relatively shallow at about 500 m depth , asŽ .expected for an off-axis site cf. Fig. 6 . Evidence of

stress control of intrusion depth is provided by theŽobservation that the measured crossover depth s sh

.s corresponds with the top of a thick intrusivev

mass that extends from 500 m to the bottom of theŽ .hole at 600 m base of intrusive not reached . There

is an apparent need to confirm the generality of theflexload structure by stress measurements in thedeepest oceanic boreholes.

1.3. Mesozoic mantle surge

1.3.1. Disruption of steady-state conÕectionThere is a wealth of evidence, from dated orogeny

and igneous activity, of a global sequence of Meso-zoic tectonic episodes that are conventionally associ-ated with a change in plate motions but can reason-ably be attributed to a surge of mantle flow, one thatapparently peaked in Cretaceous time. General re-quirements for a prolonged worldwide surge includea large source of thermal energy and a responsivemantle flow system to transmit that energy. The mostlikely energy source is the lower mantle, presumablytapped during a change from layered to whole man-

Žtle convection Machetel and Weber, 1991; Stein and.Hoffman, 1994 . A change of that kind could be

intermittent, related to inherent instability of theŽsystem Christensen and Yuen, 1985; Honda et al.,

.1993c; Weinstein, 1993 , or it could be triggered, for


example by a major asteroid impact in an oceanicŽ .region cf. Stewart et al., 1993 . The possible in-

volvement of lower mantle and mantle–core bound-ary in the Cretaceous surge of Pacific volcanism was

Ž .proposed by Larson 1991 on the grounds that thebeginning and duration of extreme volcanism coin-cided with the Cretaceous long interval of normalpolarity . Triggering of the Mesozoic surge in the

Ž .Pacific Keith, 1993 is consistent with an array ofCretaceous structures and some present-day residual

Žfeatures, including a lower mantle plume Dziewon-.ski et al., 1993 , that extends for more than 1800 km

Žacross the western equatorial Pacific Section 2, Fig..22 , a broad path that leads to the speculation that

the trigger may have been an asteroid cluster ratherthan a single asteroid impact. The apparent peak andmost persistent part of the disturbed region corre-sponds with the Polynesian superswell and SOPITAgeochemical anomaly, approximately centered onTahiti. It encompasses the Marquesas, Tuamotos and

Ž .Cook–Austral island chains Section 2, Fig. 22 .Geochemical evidence for a broad Pacific center ofdisturbance is provided by the observation that basaltsof the Polynesian superplume region have trace ele-ment signatures similar to those of the Ontong–Java

Ž .plateau Tatsumi et al., 1998 .The western extension of the disturbed region is

marked by the occurrence of abundant guyots cappedŽ .by Cretaceous reefs Hamilton, 1956; Menard, 1964 ,

Žand by anomalously thick basaltic crust Larson,1991; Ricciardi and Abbott, 1993; Abrams et al.,

.1993 , more than 25 km, for example, at the On-tong–Java plateau and about 21 km at Manihikiplateau, among the largest of the world’s igneous

Ž .provinces Schubert and Sandwell, 1989 . Radiomet-ric ages of the plateau basalt are Early Cretaceous,122 Ma for Ontong–Java, 118 Ma for a core sample

Ž . Ž .from Manihiki DSDP Site 317 Ito and Clift, 1998 .Related features include the occurrence of Jurassicsediments and pillow lava below Cretaceous layers

Ž . Ž .at ODP Site 801 188N, 1578E Larson, 1990 , aconcentric pattern of magnetic lineations and a Juras-

Žsic Aquiet zoneB Pringle et al., 1993, p. 122; cf.Larson and ODP Leg 144 Shipboard Scientific Party,

.1993 .The bordering region of the equatorial western

Pacific is marked by paired subduction zones withopposite dips. Most of the Pacific subduction zones

dip toward the adjacent continent, taken to indicatethat they are displaced convection cell boundaries atthe confluence of outflowing relatively warm sub-continental mantle and colder sub-oceanic mantle. Incontrast, the region from 148S to 198N is marked bytwo re-entrant arc systems, concave toward theoceanic side, and facing toward a focal center atabout 58N, 1758W. In addition, there are severalopposed subduction zones typically separated by a

Žcompressed wedge of detrital sediments Hamilton,.1988, Fig. 4 . The anomalous arc reversals can be

accounted for by a surge of ocean-centered mantleflow that reversed the normal sub-continent outflowand compressed pre-existing shelf and trench sedi-ments into an accretionary wedge. Massive subduc-tion and melting of those sediments is proposed tohave formed hydrous granodiorite intrusives and themajor porphyry copper deposits of the regionŽ .Solomon, 1990 .

Proposed consequences of the Pacific-centeredmantle surge include circum-Pacific orogeny, accel-erated volcanic and plutonic activity, beginning prior

Žto 165 Ma Day and Bickford, 1989; Fleming et al.,.1992 , development of inclined subduction zones

around the Pacific rim, and eastward displacement ofthe ancestral Mid-Pacific Ridge to form the EPR. Itis suggested, in accord with the proposed recyclingmodel of volcanism, that the persistence of circum-Pacific subduction long after the peak of mantlesurge is due to downward focusing of mantle flowtoward accumulated cold masses developed asresidues of repeated melting episodes. An analysis of

Ž .P-wave travel-time tomography Inoue et al., 1990identified fast anomalies, well defined at 478 to 629km depth, along the downward extensions of thePacific subduction zones. Recently identified linearfast-velocity anomalies in the lower mantle below

Ž .the Americas Grand et al., 1997 may be similarresidues, displaced deep masses that were generatedby multiple melting episodes below the ancestralMid-Pacific Ridge. Those masses have been at-

Žtributed to deeply subducted plates Van der Hilst et.al., 1997 but they are remarkably linear, not plate-

like, and residual sub-ridge masses would have therequired properties, including fast seismic velocities,

Ž .high viscosity due to depleted volatiles , and lowŽ .temperature due to depleted radioactivity . They

appear to be well defined and continuous beneath the


Americas at 800 and 1050 km depths, deepest be-neath North America.

1.3.2. Expansion and eleÕation of oceanic ridgesMesozoic expansion and elevation of mid-ocean

ridges are attributed to the proposed Mesozoic surgeŽ .of mantle flow Keith, 1993 . Related effects include

worldwide acceleration of sub-ridge subduction, andresultant widening of the MORB volcanic zone.Accelerated recycling of hydrous crustal materials atroll boundaries is proposed, in addition, to havegenerated the numerous diapirs of serpentinizedmantle in the median valley and along fracture zones,including major mantle uplifts such as the Vema and

ŽRomanche transverse ridges Bonatti and Honnorez,.1971, 1976 , the displaced Ninetyeast Ridge in the

ŽIndian Ocean Sclater and Fisher, 1974; Souriau,.1981 , and smaller mantle diapirs elsewhere, for

example along the southern wall of the Kane FZŽ .Auzende et al., 1994 , and along the Atlantis FZŽ .Blackman et al., 1998 .

The Mesozoic surge of mantle flow is proposed tohave spread out, worldwide, and eventually yielded amigrating succession of magmatic and tectonic

Ž .episodes Keith, 1993; Miller and Busby, 1995 .There are indications that the Mesozoic surge did notexpand uniformly. There is evidence, for example, ofa relatively unobstructed surge path in the southernhemisphere, delineated by the regional pattern of theADupalB geochemical anomaly in oceanic basaltŽ .Hart, 1984 and by a related pattern of low seismic

Ž .velocities in the lower mantle Castillo, 1988 .It is proposed that accelerated and relatively deep

recycling of hydrous crustal materials at roll marginsand axial convection cell boundaries led to uplift and

Ž .expansion of mid-ocean ridges see below plus amassive increase in the rate of MORB volcanism and

Žin the width of the active volcanic zone Keith,.1993 . The existence of an uplifted mid-ocean ridge

system in Late Mesozoic time, and of a broadenedzone of volcanic activity, are central to the postu-lated model of later subsidence of the ridge andnarrowing of the volcanic zone. There is good evi-dence that the North Atlantic sector of the oceanicridge system was uplifted and expanded during the

Ž .Mesozoic surge Section 3.b , but that the axis wasŽnot subject to major displacement Keith, 1993, p.

.275 . The age-related pattern of magnetic stripes is

symmetrical and remains approximately centered be-tween adjacent continental masses. The northernMAR therefore provides a simple example of theproposed ridge expansion and later subsidence.

1.3.3. Widening of the MORB Õolcanic zoneWidening of the active volcanic zone during a

major surge of mantle flow is proposed to followfrom a postulated relationship between rate of con-vective overturn and the thickness of the thermalboundary layer in uppermost mantle. A major surgeof mantle flow, such as the proposed Pacific-centeredMesozoic surge, involves a large release of thermalenergy. Dissipation of the excess heat involves in-creased thermal gradients in the upper mantle, thick-ening of the upper boundary layer and a relatedincrease in the dimensions of secondary convectiverolls within that layer. A principal effect of theenlargement of convective rolls during a surgeepisode would be deepening of roll-margin recy-cling. Under surge conditions, the roll margins, nor-mally sites of shallow recycling and related serpen-tinization, would become sites of active volcanism,thus enlarging the active volcanic zone to encompassmost of the ridge, out to the limit of initiation of

Žconvective rolls cf. Haxby and Weissel, 1986; Maiaand Diament, 1991, re volcanism along roll bound-

.aries .The proposed increase in roll dimensions and

overturn rate, and related deepening of roll-marginrecycling, presumably produced an increase in vol-ume of the low-density sub-ridge crustqmantlemixture, and consequent isostatic uplift of mid-oceanridges. The indicated Mesozoic widening of the zoneof mid-ocean ridge volcanism is attributed partly tothe increased mass of low-melting source materialŽ .sub-ridge mixture , partly to ridge uplift and conse-quent lessening of the extrusion constraint due toocean depth. Evidence for the proposed boundarylayer thickening and for enlargement of convectiverolls during the Pacific-centered mantle surge can bederived by comparing the spacing of residual rollboundaries near the Pacific center of disturbancewith the roll boundary spacing in protected areasremote from that disturbance, where convective rollspresumably have retained or regained their steady-state dimensions. As expected, the largest roll spac-ings are in the Pacific, the smallest in the protected


North Atlantic. Geoid lineations, taken to be theŽgravitational expression of roll boundaries Haxby

.and Weissel, 1986; Keith, 1993 , have typical wave-Žlengths of 55–100 km in the Atlantic Thibaud et al.,

. Ž1995 , 200 km in the Indian Ocean Cazenave et al.,

. Ž1987 , 750 to 1100 km in the Pacific Haxby, 1985;Maia and Diament, 1991; Cazenave, 1992; Smith

.and Sandwell, 1997 .

1.3.4. Limits of surge-related MAR ÕolcanismOuter limits of the zone of expanded MAR vol-

canism in Mesozoic time are marked approximatelyby two indicators of massive volcanic loading and

Ž .subsidence Fig. 8 , and, in crust of Late Cretaceousage, by an abrupt transition in volcanic topographyand an abrupt change in the variability of heat flow.North of the Charlie–Gibbs FZ, the maximal outerlimit of the expanded MAR volcanic zone was con-strained by bordering continental masses and ismarked by a thick seaward-dipping reflector se-

Ž . Ž .quence SDRS Hinz, 1981 , detected by seismicsurveys and identified as tilted seaward-thickeningmasses of basalt. The SDRS has been traced alongthe Greenland shelf and along an opposite line, morethan 2500 km long, extending from the seaward edgeof Rockall Plateau to the Lofoten Islands off the

Žcoast of Norway Mutter et al., 1984; Eldholme and.Grue, 1994 . Similar seaward-dipping reflectors have

been identified in the South Atlantic, associated withŽthe Parana and Etendeka flood basalts Gladczenko

.et al., 1997 .ŽThe drilled sections of SDRS flows Sites 918,

.989, 990 yield radiometric ages in the range 50 MaŽ .to more than 63 Ma base not reached , the youngest

equated with chron 23, of Early Eocene age. Thebasalts are compositionally similar to MORB butthose drilled at the outer feather-edge limits of the

expanded volcanic zone were extruded sub-aerially,as indicated by sections cored at DSDP sites 553 and642 on the eastern side, and at ODP site 917 on the

Ž .west Fig. 8 . Those features are superficially consis-tent with either of two contrasted hypotheses: anassociation of SDRS basalts with an early stage ofcontinental splitting or with the proposed maximalexpansion of the active volcanic zone during theMesozoic mantle surge. It appears that the basis for areasoned choice between those hypotheses mayemerge from a re-examination of the timing andlimits of the major basaltic outflows. If the Atlanticformed by continental splitting and divergent conti-nental drift, and if, as has been proposed, the sea-ward-dipping basaltic sequences represent the initial

Ž .stage of continental splitting Mutter et al., 1984 ,they should be present everywhere along the Apost-splitB continental boundaries. That does not appearto be the case for low-latitude Atlantic margins.Where the Atlantic is widest, the outer limits ofMORB volcanism at maximal expansion apparentlydid not reach the continental margins. A recentdetailed study of the Iberian continent–ocean transi-

Žtion on ODP Leg 173 Shipboard Scientific Party,.1998 noted that AThe seaward-dipping reflector se-

quences that are characteristic of volcanic marginsare nowhere seen, and on-shore synrift magmaticactivity is almost absent.B

In the equatorial Atlantic there are several signifi-cant indicators of an expanded MORB volcanic zonethat apparently peaked in Late Cretaceous time. Thefirst indicator is based on outer limits of load-in-duced tilting of the crust. Volcanic loading andsubsidence are marked by paired magnetic anomalies

Ž .33 and 34 Fig. 8 that are oppositely skewed by 408on opposite sides of the ridge axis, a skewnessdiscrepancy two to three times larger than those of

Ž .any previously observed anomalies. Cande 1978

ŽFig. 8. Northern sector of the Mid-Atlantic Ridge to show contrasted indicators of outer limits of the expanded Mesozoic ridge heavy black. Žlines , and sites that provide evidence of former shallow-water or subaerial conditions in the near-crestal region filled circles with DSDP

.numbers . To the north of Charlie–Gibbs fracture zone the landward limits of the maimal ridge and accelerated oceanic volcanism areŽ .marked by thick tilted sequences of basaltic flows: seaward-dipping seismic reflectors black abutting on continental masses. Boreholes that

Ž .intersected basalts comprising thin-edge landward limits of the expanded Mid-Atlantic Ridge green circles include DSDP 553 and 642 onŽ . Ž .the eastern side Eldholme and Grue, 1994 , and ODP 917 on the western side, off Greenland Leg 152 Shipboard Scientific Party, 1994 .

Ž .Farther south a limit of load-induced sinking of the maximum ridge is indicated by skewed magnetic anomalies 33 and 34 see text . Crestaland ridge flank sites that show evidence of former shallow-water or subaerial conditions are listed in Table 1, along with the relevant

Ž .shallowing criteria and references. Base map and magnetic anomalies are from Vogt and Tucholke 1986 .


considered the possibility that the observed skewnessreflects tectonic tilting but discounted that possibility

because it requires uniformly large tilting over anextensive region of the Atlantic. Widespread tilting


was subsequently confirmed, however, by the dis-covery of thick seaward-dipping basaltic masses on

Ž .North Atlantic margins Fig. 8 , and by ridgeward-dipping seismic reflections in crust with a range ofages. It seems evident that the occurrence of tiltedcrust over wide areas of the ocean can be attributedeither to massive volcanism during continental break

Žup and early Aseafloor spreadingB Storey et al.,.1992 or to differential loading over the proposed

broad volcanic zone of Cretaceous time. Many de-tails of the indicated oceanic subsidence remain to beworked out.

A second indicator of the outer limits of theexpanded Mesozoic zone of accelerated volcanism isprovided by an abrupt Jurassic to Early Cretaceous

Ž .change in basement relief Sundvik et al., 1984 ,Žfrom smooth volcanic topography 150 to 450 m

. Žrelief on older crust, to rough topography 450 to.120 m relief on younger crust. The authors at-

tributed the smooth–rough transition to a change

from fast to slow spreading, but the transition doesnot correspond with an isochron, as might be ex-pected from the plate tectonics model, and it is notclear why slow spreading should yield rough topog-raphy. A similar abrupt change in volcanic topogra-phy is a feature of the J-anomaly Ridge in the

Ž .western Atlantic. DSDP Site 384 Table 2 showsevidence of prolonged subaerial exposure and islocated on a prominent ridge above a steep west-fac-ing escarpment that has the character and dimensionsof a former seacliff.

A proposed third indicator of the outer limit ofexpanded Mesozoic volcanism on the MAR is anabrupt change, at a crustal age of about 60 Ma, fromrelatively uniform heat flow below marginal basinsto widely variable heat flow on the flanks of the

Ž .ridge Fig. 1 . That change is accounted for byŽdominant thermal conduction in older crust )60

.Ma and the apparent dominance, in younger crust,of heat loss by hydrothermal circulation of seawater

Table 2Limiting seismic velocities in olivine

w x w x w xaspole to 100 , bspole to 010 , cspole to 001 .Ž . ŽPropagation direction of S-waves is at 908 to the polarization or axis of vibration after Kumazawa and Anderson, 1969; McKenzie, 1979;

.Nicolas and Christensen, 1987 .


Ž .Mottl and Wheat, 1994; Booij and Staudigel, 1996 .All of the above surge indicators are consistent withthe proposed Mesozoic acceleration of volcanic load-ing and with widening of the ridge volcanic zone.

1.3.5. Contrasting features of MAR and EPRThe northern MAR is used, above, as the prime

example of ridge dynamics because there is evidencethat it was relatively protected from disruption by theproposed mantle surge of Mesozoic times and conse-quently represents an approach to a steady-state con-dition. Characteristic indicators of the relatively pris-tine character and prolonged existence of the NorthAtlantic ridge include its mid-oceanic position,equidistant from major shields, the symmetrical dis-position of magnetic anomaly ages and geophysicalparameters, and its association with a large positivegeoid anomaly, interpreted to imply strong sub-axialdownflow, focused by an accumulated mass of cold

Ž .and viscous residue Fig. 4 , relatively depleted inradioactive elements and volatiles as a result ofprolonged sub-axial recycling and melting.

In contrast, the EPR is taken to be the mostseriously disrupted segment of the oceanic ridge, aconvection cell boundary that was displaced east-ward from a former mid-Pacific position and is nowcharacterized by pronounced asymmetry of magnetic

Žages and geophysical parameters MELT Seismic.Team, 1998 . Most notably, the relocated ridge ap-

parently has been separated from most of its accumu-Žlated cold residue presumed to have been a feature

.of the ancestral Mid-Pacific Ridge , and as a resultthe EPR lacks the indications of strongly focused

Žsub-axial downflow median valleys and a broad.positive geoid anomaly that characterize the north-

ern MAR. It seems likely that the linear high-veloc-ity anomalies recently identified in the lower mantle

Ž .below the Americas Van der Hilst et al., 1997 maybe displaced sub-ridge residues from the ancestralMid-Pacific Ridge.

Despite the above significant differences from theMAR model, the detailed geophysical studies of theEPR MELT experiment yield important insights re-

Ž .garding ridge dynamics Forsyth et al., 1998 . Theoccurrence of sedimentary layers capping tiltedmarginal blocks adjacent to the Juan de Fuca seg-

Ž .ment of the EPR Davis and Lister, 1977 is consis-tent with the proposal that the major sub-axial con-

vection cell boundary of the EPR is a displacedboundary, relict from an ancestral Mid-Pacific Ridge,and that it moved eastward into a near-continentregion formerly occupied by undistorted, nearlyflat-lying layers of terrigenous sediment.

Recent seismic surveys across the EPR provideevidence of a broad sub-ridge LVZ previously

Ž .thought to be a melt zone Forsyth et al., 1998 . Inthe proposed recycling model of MORB volcanism,it is identified primarily as a zone of hydrationreactions, where crustal materials are carried down at

Žfracture zone boundaries proposed margins of con-.vective rolls , and the released water is taken up in

the exothermic serpentinization of mantle peridotite.The observed increase in shallow-path P- and S-wave

Ždelays, on approach to the axis Toomey et al., 1998,.Fig. 1 , can be attributed to the progressive addition

of recycled hydrous material to uppermost mantlealong the lines of fracture zone subduction.

A separate study of distant-source P-waves arriv-Ž .ing on steep paths Toomey et al., 1998, Fig. 4

showed that those with near-vertical paths near theEPR axis are relatively fast. Except for a narrowcentral region of relatively slow seismic velocityŽ .axial subduction of crust? , the velocity of steep-pathP-waves decreases progressively with distance fromthe axis, and reaches average values at about 100 kmout. The results are similar to those of an earlier

Ž .P-wave transect of the MAR Blackman et al., 1993and are consistent with sub-axial downwelling rather

Ž .than an upwelling plume Section 1.1.12ŽIt seems likely that the MELT seismic team For-

.syth et al., 1998 have located a principal candidatefor the AtrueB zone of magma generation, a relativelysmall sub-axial zone at about 75 km depth, whichthey tentatively identified as a zone of anomalous

wmelting of an Aembedded heterogeneityB op. cit., p.x1217 and Fig. 3 . A narrow sub-axial melting zone is

consistent with the proposed recycling model ofMORB magma generation and avoids the difficultproblem of trying to account for long distance lateralmelt migration from a broad plume into a narrowaxial volcanic zone. The limits of the deep zoneattributed to magma generation are not well con-strained, but it is shown as offset about 50 km to thewest of the ridge axis, and that offset, along withother asymmetric features of the EPR, favors theconcept of a displaced ridge which reached maximal


displacement during the Mesozoic mantle surge, andis beginning to return toward a mid-oceanic position,presumed to be the steady-state geometry of thePacific ridge.

1.3.6. Subaerial exposure and subsequent sinking ofthe ridge

The postulated former elevation and sub-aerialexposure of the mid-ocean ridge system is supportedby the evidence, along ridge crestal regions, of shal-low-water deposition, including wave-worn basalticcobbles and sand, weathered and oxidized basalt,some capped by oxidized sediment or by shallow-

Ž .water reefal and lagoonal limestone Keith, 1993 . InŽ .the North Atlantic examples Table 1 , several of the

sites of formerly emergent or shallow-water condi-tions are on transverse ridges, for example the 500-m-thick platform of reefal and lagoonal limestone

Žthat caps the Vema transverse ridge Bonatti et al.,.1994 , but most of the critical sites are on or close to

Ž .the ridge axis Fig. 8 . Approximate post-extrusionsubsidence, as indicated by the present sub-sea depthsof formerly emergent or shallow-water layers, ranges

Žfrom 922 m for the most northerly site DSDP Site.409 , to more than 4000 m at site 556.

Further evidence of former ridge exposure is theextensive denudation of oceanic crust, developmentof deep canyons and trellis drainage patterns along

Ž .fault scarps of the MAR Tucholke et al., 1997 . Theauthors attributed the modified topography to sub-

wocean mass wasting but the erosional features op.xcit., Fig. 3 , out to about 300 km from the axis, favor

Recent sub-aerial exposure and erosion of the ridgecrest. Former broader sub-aerial exposure, and pro-gressive subsidence, is indicated by borehole inter-sections at off-axis sites, for example at DSDP sites407 and 384. Exposure history of Reykjanes Ridge isindicated by former seacliffs, now submerged, and

by the stratigraphic evidence of three DSDP bore-holes drilled along a ridge-normal transect at 62–

Ž .648N Fig. 9 . The youngest seacliff is at the outerŽ .limit of the Icelandic shelf Saemundsson, 1986 and

the older ones are indicated by two V-shaped traces,Ž .visible on the SEASAT gravity map Haxby, 1985 ,

that extend southward along Reykjanes Ridge toabout 900 km from the present shore of IcelandŽ .Keith, 1993 . The symmetrical V-shaped traces werepreviously attributed to southward flow of astheno-sphere from a postulated Icelandic mantle plumeŽ .Vogt, 1971; White et al., 1995 but seismic reflec-

Žtion profiles across the ridge Talwani et al., 1971, p..490 show that the traces have the aspect of sea

cliffs, and their 600 m height is similar to that ofŽcliffs on the present shores of Iceland Keith, 1993,

.p. 222 . One of the seacliffs shows up on a seismicprofile along the ridge-normal transect through DSDP

Ž .holes 407–409 Keith, 1993, p. 214 . The stratigra-phy of those borehole sections provides evidence ofbroad exposure of the ridge followed by ridge subsi-dence, narrowing of the active volcanic zone and

Ž .migration of the seacliff Fig. 9 . The deduced suc-Ž .cession of events below includes citations to rele-

Žvant pages in the DSDP report Shipboard Scientific.Party, 1978b .

At site 407, 300 km from the ridge axis and nowsubmerged to 2.4 km depth, the lowermost drilled

wbasalt is an oxidized vesicular pillow lava op. cit., p.x39 , characteristic of a shallow water extrusion site.

Intermittent sub-aerial exposure is confirmed by theoccurrence of basalt cobbles with thick iron oxide

w xweathering rinds op. cit., p. 73 . The indicated widthŽ .of exposed MAR at that time 36 Ma is about 600

km. The final stage of volcanism was in shallowwater, but there are indications of rapid post-volcanicdeepening. The uppermost basalt is vesicular and theoverlying sediment contains a glauconitic layer andoccasional turbidites, indicators of shallow water

Fig. 9. Selected AsnapshotsB of a Cretaceous to Recent time sequence for a sector of the subsiding Mid-Atlantic Ridge, Azores toCharlie–Gibbs fracture zone, to show narrowing of the active volcanic zone, the proposed mechanism for generating oceanic magneticstripes. The narrowing blue-green color pattern shows the extent of active MORB volcanism at each selected stage of ridge subsidence andcooling. The trailing edge of the crestward-retreating outer limit of volcanism and mid-crust gabbro intrusion is taken to be associated withprogressive cooling of the ridge and with the migrating trace of the blocking temperature isotherm, at which remanent magnetism is frozen

Žin, thus recording inclination and reversals of the Earth’s magnetic field. Approximate magnetic anomaly ages upper right corner of each. Ž . Ž .panel are from Kent and Gradstein 1986 cf. Harland et al., 1990 .


deposition. Those beds are followed by more than200 m of deep-water nannofossil ooze. Long term

proximity of the seacliff is suggested, based on theabundant occurrence of interflow basalt breccia


through a 100-m section of submarine lava flowsw xop. cit., p. 39 , and of both rounded and angular

wbasalt fragments in overlying sediments op. cit., pp.x21, 37, 40, and 45 . A similar shallow to deep

extrusion sequence is evident at site 408, although itlacks the indications of intermittent sub-aerial expo-sure. A particularly significant test could be made bydeepening hole 408. The proposed subsidence modelleads one to predict that a deepened hole shouldencounter weathered sub-aerial lavas below the shal-low-water sequence. By the end of Miocene timeŽ .ca. 24 Ma , sites 407 and 408 were both volcani-

Žcally inactive and deeply submerged accumulating.fine-grained pelagic sediment .

In contrast with sites 407 and 408, site 409, in thehigh near-axial region, was a site of continuoussub-aerial volcanism, a condition that apparently per-sisted until the final stage, as indicated by the contin-ued accumulation of frothy basalt, to 38% vesicular-

w xity op. cit., pp. 161, 168 , and the absence ofw xseawater alteration op. cit., p. 178 . It seems appro-

priate to pose the question as whether some of thefrothy basalt may be sub-glacial, extruded beneath aPleistocene ice sheet. Site 409 evidently continuedsinking, along with the whole ridge. It eventuallywas covered by 80 m of Upper Pliocene and Pleis-tocene glacio-marine sandy muds containing gradedturbidites and coarse glacial erratics, abundant inuppermost Pleistocene.

The above outline can be compared with a platetectonics interpretation of the same ridge transect, as

Ž .proposed by Duffield 1978 . He calculated the aver-age vesicularity of the cored basalt as: hole 407.6%,hole 408.13%, and hole 409.27%, and attributed thevesicularity change to a model involving rise of theridge and progressive shallowing of water depth atthe postulated axial source of oceanic crust. TheDuffield model is generally consistent with thestratigraphy of holes 408 and 409, but cannot ac-count for hole 407 stratigraphy, including the vesicu-lar weathered basalt at the base and the prolongedavailability of angular basalt breccia, here attributedto proximity of the seacliff.

Indications of an earlier, even broader expanse ofsubaerial exposure of the ridge are provided bystratigraphic evidence from DSDP Site 384 in thewestern Atlantic, at 40.38N, 51.68W, in water depth

Ž .3909 m Shipboard Scientific Party, 1976 . Borehole

384 penetrated a continuous sedimentary sequenceacross the Cretaceous–Tertiary boundary and inter-sected a shallow-water, rudist reef limestone, of

Ž .early Cretaceous Aptian age, that shows evidenceof subaerial exposure. The Cretaceous limestones areunderlain by highly vesicular weathered basalt at-tributed to shallow water or subaerial extrusion. Site384 is about 850 km from the axis of the MAR, andthus provides an estimate of about 1700 km for theexposed width of the active MAR volcanic zone inEarly Cretaceous time. There is evidence of evenmore extensive ridge exposure in the Pacific, notablythe shallow-water Cretaceous reefs and karstic to-

Žpography on western Pacific guyots Menard and.Chase, 1970; Winterer, 1998 .

The proposed former exposure of Reykjanes Ridgeand of some other ridge segments is supported bygeochemical indicators of recycling of sub-aerial

Ž .basalt to the MORB source Section 2 . In additionŽ .to indicators of weathering high Fe content and

Ž 3 4 .cosmic ray exposure high Her He ratios , Reyk-janes Ridge samples exhibit a northward decrease in

Ž .Na O content Langmuir et al., 1992, p. 204 at-2

tributable to a decrease in the proportion of seawateraltered sub-oceanic basalt in the recycled crustalmaterial. Similar variability in the proportions ofrecycled sub-aerial and sub-oceanic basalts is indi-cated by a negative correlation of Fe content with Na

Žand K in basalts from the EPR Langmuir et al.,.1992, figs. 25c and 28 .

1.3.7. Volcanic zone contraction and the generationof magnetic stripes

An essential feature of the proposed ocean ridgesystem is that all of the abnormal features thatresulted from the Mesozoic surge of mantle flow:expansion and thickening of the LVZ, uplift of theridge, accelerated volcanism, broadening of the ac-tive volcanic zone, were subject to gradual Mesozoicto Recent relaxation and retreat, toward a steady-statesystem, a predictable effect of the slowing of convec-tive overturn and volcanism, and the gradual dimin-ishing, via return-flow gyres, of the accumulatedlarge volume of sub-ridge subduction mixtures. Theage-denominated sequence of magnetic anomalies,conventionally attributed to sea floor spreading, isproposed to result, instead, from gradual narrowingof the active volcanic zone. The postulated post-


Cretaceous reduction in boundary layer thicknessand roll dimensions, is evidenced by large differ-

Ž .ences in the spacing of fracture zones roll margins ,from as much as 1500 km in the central Pacific,massively disturbed by the Mesozoic mantle surge,to about 55 km in the protected North Atlantic.

Roll margin recycling is proposed to be the prin-cipal mechanism of recycling of hydrous crustalmaterials, with resultant serpentinization of uppermantle peridotite and development of the sub-ridgeLVZ. Under present conditions, roll margin recy-cling presumably is not deep enough to yield sub-stantial partial melting and basalt generation. It isproposed, however, that deeper roll-margin recy-cling, associated with the larger rolls of the Meso-zoic mantle surge, would have yielded a broad zoneof volatile-promoted MORB volcanism.

Two secondary effects are attributed to ridge sub-Ž .sidence: 1 enhanced crustal cooling due to the

change from sub-aerial to sub-oceanic volcanism,Ž .and associated seawater circulation, and 2 the crest-

ward migration of a limiting ocean depth that con-

strains magmatic injection and magmatic bulging ofoceanic crust and thus shuts off the Avolcanic valveB,the mechanism of bulge-crest fissuring and extru-sion. An effective limit to sub-ocean extrusion undermagmatic pressure is related to the positive slope ofthe basalt liquidus in P–T space. A principal result,at near-liquidus pressure and temperature, is that asignificant increase of pressure will cause the magma

Ž .to crystallize Keith, 1993 .ŽNarrowing of the zone of MORB volcanism Fig.

.10 and crestward cooling of the ridge will be ac-companied by crestward migration of the blocking

Ž .temperature isotherm Fig. 11 , at which the mag-netic minerals acquire remanent magnetism and thusrecord inclinations and reversals of the Earth’s mag-netic field. Differences in the length scale of themagnetic stripe pattern, currently taken to implyvariable spreading rate, are attributed, instead, tovariable rates of narrowing of the volcanic zone.Both the spreading model and the proposed coolingmodel of magnetic stripe generation are faced withthe nagging problem of reconciling the concept of a

Ž .Fig. 10. Idealized succession of diagrams of one side of the mid-ocean ridge axis at A , to illustrate the development of oceanic magneticstripes, Mid-Cretaceous to Present, as a result of gradual narrowing of the active MORB volcanic zone, cooling of the crust and crestward

Ž .migration of the blocking temperature isotherm B , at which the magnetic minerals acquire remanent magnetism and thus record the Earth’sfield. The gray pattern to the right of B indicates the portion of the ridge for which the principal magnetic source remains above blockingtemperature. The time-frame AsnapshotsB of magnetic stripe development are arbitrary, selected to match those in Fig. 9. Geomagnetic

Ž .polarity and geologic time scale are from Kent and Gradstein 1986 .


simple magnetic source of linear anomalies with theabsence of such a source in the drilled upper 600 m

Ž .of oceanic crust Hall and Robinson, 1979 . Theproblem apparently can be resolved by evidence thatthe dominant magnetic anomaly source lies at deeper

levels, below the uppermost crustal layers reached bydrilling. Recent results favor the hypothesis thatthere is a significant contribution from secondary

Žmagnetite, in basal layer-2, in layer-3 gabbros VonHerzen et al., 1991, p. 555; Kikawa and Ozawa,


.1992; Pariso and Johnson, 1993 or in serpentinizedŽupper mantle peridotite Toft and Arkani-Hamed,

. Ž .1992 . Tivey and Tucholke 1998 made a sea-surfacemagnetic survey over a near-axial segment of theMAR and concluded that magnetization of the lowercrust probably is the dominant source for off-axismagnetic anomalies.

Supportive evidence for a deep magnetic anomalysource can be derived from the relationship betweencrustal structure and the skewness of magnetic

Ž .anomalies. It was reported Cande, 1978 that, withthe exception of the unusual Cretaceous anomalies33 and 34, attributed to outer limits of Mesozoic

Ž .volcanic loading Fig. 8 , most of the marine mag-netic anomalies are not significantly skewed, even incases where tectonic tilting is evident in drill coresections or seismic reflection profiles. The apparentconflict can be resolved by assuming that the entirecrust was tilted and that the magnetic anomalies aredue mainly to deeper rock masses that were magne-tized during delayed, post-tilt cooling through the

Ž .blocking temperature Section 1.2.4 .It will be evident that the spreading model of

Žplate tectonics and the proposed cooling model Fig..11 will yield similar sequences of age-related mag-

Žnetic anomalies Kent and Gradstein, 1986; Harland.et al., 1990 . The cooling model is independent,

however, of several questions and anomalies associ-ated with the spreading model. Unaccounted for bythat model are the puzzling occurrences of reversely

Žmagnetized rocks in the Median Valley Macdonald,.1982; Palmason, 1986, p. 93 , and the failure of

detailed studies, for example in the FAMOUS areaof the MAR, to confirm either symmetrical spreadingor the existence of a narrow zone of crustal accre-

tion, 1–2 km wide, that could remain stable overtime and yield the observed pattern of magnetic

Žstripes Macdonald and Luyendyk, 1977; Saemunds-.son, in Vogt and Tucholke, 1986 . Goldstein et al.

Ž .1994 showed, by uranium series dating of basaltson the EPR, that active volcanism is not restricted toa narrow near-axial zone. They found young basaltsthat had been erupted as much as 4 km off axis. Offaxis volcanism, and the prior occurrence of a muchwider zone of MORB volcanism is consistent withstratigraphic evidence of DSDP boreholes 407 to 409drilled along a ridge-normal transect of Reykjanes

Ž .Ridge Fig. 9 .

1.3.8. Ages of sediment caps and interlayers on theridge

One of the lines of evidence that has been takento confirm the age sequence defined by the spreadingmodel of magnetic stripe formation is the existenceof a parallel and matching sequence of ages ofAbasalB sedimentary layers that lie directly on theuppermost basaltic crust. That coincidence is notdiagnostic, however, because the same matching se-quence of magnetic anomaly ages and sediment ageswill result from the proposed ridge-cooling model ofmagnetic stripe generation. The development ofmatching ages can be understood by reference to Fig.9, in which narrowing of the active volcanic zoneŽ .green pattern is represented by several steps andsuperimposed on the sequence of numbered mag-netic anomalies. At 55 Ma, for example, magneticanomaly 24 of that age will recently have formed bycooling of the crust through the blocking tempera-ture, and lower Eocene sediments of the same age

Fig. 11. Simplified sequence of sections representing the subsidence and cooling of Reykjanes Ridge, based on data from DSDP Sites 407,Ž . Ž .408 and 409 62–648N Shipboard Scientific Party, 1978b, pp. 39, 117, 126, and 167 , to show consistency between borehole stratigraphy

Žand the proposed model for generating oceanic magnetic stripes by ridge cooling following a Cretaceous peak in volcanic activity see text. Ž .and Fig. 10 . The three early stages of cooling and migration of blocking temperature, B , represent conditions at magnetic anomaly ages

36, 24, and 6 Ma, chosen to approximate the stratigraphic boundaries between Eocene, Oligocene, Miocene and Pliocene. The 600 mŽ . Žseacliff now submerged between sites 408 and 409 is shown on DSDP seismic reflection profiles op. cit., Shipboard Scientific Party, p.

.164; cf. Talwani et al., 1971; Keith, 1993, p. 214 . The rate of erosional retreat of the seacliff during exposure is arbitrarily set at 1mrcentury.

Ž .Miocene sediments and the Pliocene of holes 407, 408 are siliceous nannofossil chalk and ooze, glauconitic near the basaltic base ofsection 408, and containing both rounded and angular basalt fragments above the upper basalt of section 407. The Pleistocene, the UpperPliocene of Hole 409, and the Oligocene between upper and lower basalt of section 407, are calcareous sandy mud intermixed withturbidites. All have variable proportions of volcanic ash. Sediments are represented by gray shades in the stratigraphic section. Hot crustabove blocking temperature is represented by a dark gray pattern in the time sequence sections.


will have been deposited on overlying young basalt,recently formed at the retreating outer limit of activevolcanism. Lower Eocene sediments will also be

Ž .deposited on the near-crestal zone green pattern butthe sediment volume decreases progressively withdistance from continental sources and the lowerEocene sediment will occur mainly as thin layersintercalated within volcanic flows of the still activevolcanic zone.

It is evident that a similar matching sequence ofmagnetic anomaly ages and overlying sediment ageswill develop as the volcanic zone continues to nar-row and the crust cools through the blocking temper-ature. It is suggested that a reasoned choice betweenthe spreading model and the Avolcanic narrowingqridge coolingB model can be based on a series of testboreholes through upper crust on the crestward sideof a well-preserved sedimentary layer. In our chosen

Ž .example 55 Ma crust and magnetic anomaly , thespreading model leads one to expect a cap of younger

Ž .sediments no Lower Eocene on the near-axial side,while the cooling model leads to the expectation thatthere will be thin interlayers of Lower Eocene sedi-ment in near-crestal basaltic crust as well as youngersediment interlayers in uppermost crust. Some sedi-mentary interlayers are sufficiently lithified to be

Ž .cored see below . The proposed test presumablywill require improved methods for recovering uncon-solidated sedimentary interlayers and fossils, nownormally lost in the process of drilling and corerecovery.

The spreading model involves a simple, single-stage process of crustal generation and spreadingfrom a narrow axial zone; it does not postulatesubsidence or long-term volcanism and sedimenta-tion at the source. Consequently it does not lead oneto expect sediment interlayers, except at very shal-low depth in the crust. On the other hand, the

Ž .proposed flexload model of ridge structure Fig. 6and magnetic anomaly generation by cooling of themaximum ridge, leads to the prediction that sedimentinterlayers should be a common feature throughoutthe crust, developed by long-continued sedimentationalong with volcanism and subsidence within thegradually narrowing MORB volcanic zone of Creta-ceous to Recent times. The occurrence of sedimentinterlayers at various depths in borehole sectionsthus provides a key point of evidence, confirming the

proposed ridge cooling model. The following exam-ples are selected to show depth of sediment interlay-ers in borehole sections at a wide range of distancesfrom the axis of the MAR. Cited depths are relativeto AbasementB, the top of the uppermost basalt.

Hole 412 was drilled at a near-axial site on theMAR, at magnetic anomaly age of about 1.6 Ma.Basement consists of phyric basalt flows interlayeredwith limestone that shows bedding dips up to 308,

Ž .consistent with the proposed flexload model Fig. 6 .Nannofossil chalk interlayers were cored as deep as

Ž129 m sub-basement Shipboard Scientific Party,.1978b, v. 49, p. 339 . Hole 395 was drilled to the

west of the MAR axis within normal polarity mag-netic anomaly a4. That hole intersected two brecciaswith a clay matrix and two carbonate-cemented brec-cias, one at 184 m sub-basement, containing traces

Žof microfossils Shipboard Scientific Party, 1978a, v..45, p. 131 . Hole 332, drilled about 30 km west of

the MAR axis, intersected massive to pillowed basaltwith numerous sediment interlayers that comprise asmuch as 70% of the upper section. Some chalkinterlayers occur as deep as 544 m sub-basementŽ .Shipboard Scientific Party, 1977, p. 15 . Hole 417was drilled on the western outer flank of the MAR inoceanic crust of Cretaceous age. Several limestoneinterlayers were intersected, including a 15-cm-thick

Žlayer at about 62 m sub-basement Shipboard Scien-tific Party, 1979, pp. 25, 75; Hall and Robinson,

.1979 .Ž .The array of evidence above is taken to favor an

Earth model characterized by sub-continent up-welling and sub-ocean convergent flow, toward ma-jor convection cell boundaries at the axis of themid-ocean ridge system. In contrast with the platetectonics model, in which the marine magneticanomaly pattern is attributed to ocean-floor spread-

Žing and taken as a measure of continental drift e.g.see Klitgord and Schouten, 1986; Srivastava and

.Topscott, 1986; Van Andel, 1994 , the proposedridge-cooling model of magnetic stripe formationdoes not involve spreading or continental splitting,and there are no implications regarding continentaldrift, except that the continents are presumed to befloating in the mantle, each focusing one or moreupwelling plumes, and free to move in response tochanges in the global convection pattern. The princi-pal continental drift will be related to mantle surge


episodes, and there will be a strong tendency forcontinents to re-establish their separate positions aspart of a return to a steady-state flow regime.

The above comparison of alternative dynamicmodels is focused on a key question: spreadingversus convergence of sub-ocean mantle flow. Read-ers may wish to use the table below to list their owninterpretation of the relevant clues.

A reader’s scorecard about the direction of sub-oceanmantle flow

Clues Spreading Convergence

1.2.3.4,5.6.7.8.9.10.11.12.13.14.15.

2. MORB recycling, cold spots and oceanic is-lands

MORB magma is conventionally attributed to de-compression melting in sub-axial plumes of mantle

Ž .peridotite Morgan et al., 1992 , oceanic island basaltŽ .OIB to melting of deep source plumes originatingat the core mantle boundary. The plume model fails,however, to account for the observed regional varia-tions in trace element and isotopic signatures ofMORB and OIB. In addition, it is faced with several

Ž .puzzling questions: 1 How is it possible to collect adisseminated melt fraction into a narrow central zoneof magma supply, in view of evidence that mantleplumes have large dimensions and that the meltfraction will be widely disseminated on grain bound-

Ž .aries? 2 How can the upper mantle continue to

yield basalt by decompression melting if, as we mustassume, it has been thoroughly depleted by persistentand episodic partial melting since earliest Archean

Ž . Ž .times Stolper et al., 1996b ? 3 How can the up-welling model account for the abrupt oceanic tocontinental shift from monotonous MORB in theGulf of Aden to the wide variety of igneous rocksŽrhyolites, phonolites and carbonatites, as well as

. Ž .basalts that are associated with the African rifts? 4In view of the high temperatures required for drymelting of a peridotite plume, how can one accountfor the lack of an associated broad heat flow anomalyover the Hawaiian swell and other active volcanicsystems?

It is suggested that the answers to the abovequestions emerge from a proposed model of basaltgeneration by heterogeneous melting at sites wherevolatile-enriched crustal materials are recycled intothe upper mantle. The proposed model is based on an

Ž .array of evidence Section 1 that is contrary to theplate tectonics model and favors focused upwellingbeneath ancient continental shields, focused down-flow below mid-ocean ridge axial zones and below

Ž .oceanic islands cf. Shaw and Jackson, 1973 . Arecycling model is widely accepted for the trench–arcsystems, and has been proposed in general terms forMORB, without identification of subduction sitesŽsee Kurz et al., 1982a,b and earlier investigations

.cited by them . It will be evident that remote subduc-tion at recognized Indonesian and circum-Pacificsubduction zones cannot account for the observedregional correlations between compositional varia-tions of MORB and the isotopic signatures of local

Ž .sediments see below . It is proposed that heteroge-neous recycling mixtures accumulate at convection

Ž .cell boundaries roll boundaries and axial zone , andthat new basaltic magmas are generated by volatile-promoted partial melting in sub-ridge crust–mantlemixtures in which recycled basaltic crust and interca-lated sediments are dominant components. There is,presumably, some selectivity in the recycling pro-cess, for example a tendency for resurgent basalticintrusion and extrusion at the ridge axis, in contrastwith a tendency for more viscous granitoid melts tobe blocked by water-cooled oceanic crust, and even-tually to be incorporated in a return-flow gyre, sub-ocean to sub-continent, a source of continental re-generation.


2.1. A recycling model for genesis of MORB

2.1.1. Oceanic LVZ and the MORB sourceEvidence of heterogeneous sources of oceanic

lavas is provided by variable major element compo-Ž .sition Perfit, 1999 , and by extreme trace-element

and isotopic variability, for example by a factor of70 for Ba concentration in a restricted 10=40 km

Ž .area of the EPR Prinzhofer et al., 1989 . Furtherevidence is provided by the common overenrichmentof magmaphile elements in MORB, well above theconcentrations predicted to result from fractional

Ž .crystallization Nielsen et al., 1995 , and by thecoexistence of depleted and enriched primary magmapreserved as melt inclusions in MORB olivine and

Žspinel Sobolev and Shimizu, 1994; Kamenetsky,.1996 .

Mantle heterogeneity is further indicated by ex-treme isotopic variability, for example by a wide

87 86 Ž .range of Srr Sr Fig. 13 , including a range from0.7029 to 0.7040 within a single borehole at Site 561Ž .Bryan and Frey, in Vogt and Tucholke, 1986 .Analyses of Pb, Sr and Nd isotope ratios on the same

Žsample suites Cohen et al., 1980; Dupre and Alle-.gre, 1983; Hofmann, 1997 also showed significant

departures from bulk Earth isochrons, taken as evi-dence against a uniform mantle source of oceanicbasalt. The Atlantic data sets show a positive correla-tion of 87Srr86Sr with 206 Pbr204 Pb and a negative

143 144 Ž .correlation with Ndr Nd. Cohen et al. 1980suggested, as one possibility, a previously proposedmodel involving recycling of continent-derived sedi-ments into the MORB source.

A study of variable trace element ratios in MORBŽ .Jacobsen and McDonough, 1996 and of the Nd–Hf

Ž .isotopic array Lewis and Smith, 1997 confirmedthe proposed recycling models. A recent review of

Ž .Sr, Nd, and Pb isotope ratios Hofmann, 1997 andanalyses of 21 incompatible trace elements in ninehand-picked samples of MORB glass from the In-

Ž .dian Ocean Ridge Rehkamper and Hoffman, 1997led the authors to attribute the compositional range,and differences from Atlantic MORB, to contamina-tion with a heterogeneous mix of recycled oceaniccrust plus small variable amounts of old sediment. Arelatively large sedimentary contribution to IndianOcean MORB is consistent with the abundant sedi-ment supply from the Himalayan Range. The pro-

posed model incorporates the above suggestions intoa general recycling model that includes old oceaniccrust plus pelagic and terrigenous sediment.

Indications of recycling to the MORB source leadto a re-examination of the sites of subduction andrecycling. It seems apparent that a mid-ocean crustalreservoir and focused local recycling are required. Itis suggested that the first stage of recycling occurs atconvective roll margins, marked by fracture zones,and that roll boundary recycling produces a crust–mantle mixture that is subject to crustal dehydrationand mantle hydration reactions and is detected as the

Žseismically defined sub-ridge LVZ fine dot pattern.of Fig. 4 , and as a zone of enhanced electrical

Ž .conductivity Karato, 1990 . The identification ofŽfracture zones as roll margins Haxby and Weissel,

.1986; Robinson et al., 1988 , and as sites of shallowŽ .recycling of hydrous crustal materials Keith, 1993 ,

is consistent with the associated diapirs of serpenti-Žnite and serpentinized mantle peridotite Bonatti and

Honnorez, 1976; Fox and Gallo, in Vogt and Tu-.cholke, 1986; Cannat, 1993 , presumably formed by

hydration of mantle peridotite at uppermost levels ofrecycling and mixing. The evidence favors the hy-pothesis that the subridge LVZ is a zone of hydra-tion, not a zone of partial melting. A recycling originof the oceanic LVZ is consistent with indications of

Žefficient mixing by toroidal flow Kellogg and.O’Connell, 1998 , and with evidence that the sub-

ridge LVZ is most clearly defined in the upperŽ100–150 km Su et al., 1992; Zhang and Tanimoto,

.1993, 1994; Montagner and Anderson, 1989 .Several investigators have proposed a zone of low

viscosity, corresponding to the seismic LVZ. Robin-Ž .son et al. 1988 concluded, from a numerical model

of small scale convection in upper mantle, that thetypical 150–250 km wavelengths of Atlantic geoidlineations, interpreted as roll boundaries, indicate alow-viscosity boundary layer 75–125 km thickŽ .0.5l .

The sub-ridge LVZ is conventionally attributed toelevated temperature and partial melting. Experimen-tal measurements have shown, however, that there isa gradual sub-solidus reduction of seismic velocitiesŽ .and increased attenuation , but that there is noabrupt change of either at the onset of partial meltingŽ . Ž .Sato et al., 1989 . Jung and Karato 1997 chal-lenged the common belief that low seismic velocities


and high attenuation are caused by the presence of aŽsmall melt fraction. They pointed out Karato and

.Jung, 1998 that a small melt fraction in peridotitetends to extract the available water from surroundingminerals such as olivine and may produce an in-crease of seismic velocity. The available data favorthe proposition that the shallow level sub-oceanic

ŽLVZ can be attributed to premelting effects Richet.et al., 1994 and to hydrous conditions rather than to

Ž .pervasive partial melting Popp and Kern, 1993 .Ž .The measurements of Christensen 1966 showed

that serpentinization of mantle peridotite can accountfor reduced seismic velocities similar to those ob-served in the sub-ridge LVZ. In addition, the releaseof hydrogen, a by-product of the serpentine-formingreaction, has been shown to produce a significantincrease in the electrical conductivity of olivineŽ .Karato, 1990 , a reported characteristic of sub-ridgeLVZs.

The thickness and extent of the sub-ridge LVZhave been defined approximately by S-wave tomog-

Ž .raphy. Woodhouse and Dziewonski 1989 and ZhangŽ .and Tanimoto 1993 showed that all ridges are

underlain by broad LVZs, with slowest velocities inthe upper 100–200 km. Some investigators havededuced deeper LVZ limits to 300–400 km in someregions, for example below the Gulf of Californiaand below sites of accelerated volcanism, such as

Ž .Tristan, Azores and Iceland Vasco et al., 1994 .An essential feature of the proposed model is that

oceanic crust, and the crustqmantle mixture of thesubridge LVZ, are the principal materials being recy-cled to the sub-axial source of MORB. The proposedequivalence of MORB source material with the sub-ridge LVZ is evidenced by an observed associationof basalt chemistry with seismic S-wave velocitiesŽ .Humler et al., 1993 , and a North Atlantic associa-tion, strongest at 150 km depth, between the depthextent of the LVZ and several geochemical anoma-lies, notably the positive anomalies of FeO andCaOrAl O , and a negative anomaly of Na O, in2 3 2

basalt from the Azores and Reykjanes–Iceland re-Ž .gions Zhang et al., 1994 . All of those anomalies

can be accounted for by recycling of basalt depositedin sub-aerial and shallow-water environments of theoceanic islands: indicators include enhanced cosmo-

Ž .genic helium-3 Sections 6.08 and 7.03 , FeO fromŽ .weathered oxidized basalt Fig. 13 , elevated CaO

from intercalated shallow-water carbonates, depletedNa O representing sub-aerial basalt that is free of2

seawater alteration.An association with recycling is further indicated

by the observation that the LVZ is most clearlydefined by minimal S-wave velocities in regionsproximal to sources of terrigenous sediments, e.g. innarrowest Atlantic near Iceland, in northern, near-

Žcontinental EPR and Gulf of California Woodward.and Schnieder, 1993 , and in the northern Indian

Ocean, a region of abundant detrital sediment fromŽ .the Himalayas Inoue et al., 1990 . The latter associ-

ation is supported by the conclusion of HofmannŽ .1997 that recycled continent-derived sediment is amore significant component in Indian Ocean MORBthan in MORB from either the Atlantic or the Pa-cific.

2.1.2. Volatile components of MORBThe evidence, above, for recycling of oceanic

crust and sedimentary components is supported byindications of recycling of atmospheric componentsand seawater to the MORB source. Krypton andxenon entrapped in vesicles and dissolved in rapidlychilled glassy margins of MORB are characterizedby isotope ratios essentially the same as those of

Žatmospheric gases Sarda et al., 1985; Jambon et al.,.1985; Fisher, 1994a,b . The dominant pathway of

atmospheric contamination is considered in Sections2.2.4 and 2.4.4.

Excess chlorine in MORB and OIB can be at-tributed to recycling of seawater-altered basalt, aprocess that is effectively similar to the previouslyproposed assimilation of seawater-influenced wall

Ž .rocks Michael and Schilling, 1989 . The volatilecomponents of oceanic basalt, including methane andcarbon dioxide, can be attributed, in part, to sedi-ment recycling. Abundant methane at outgassing sitesalong the MAR was attributed to serpentinization of

Ž .mantle peridotite Charlou et al., 1998 , whereasmethane in Indian Ocean gabbro was attributed to

Ž .recycled organic matter Kelley et al., 1994 .Carbon dioxide is the dominant volatile compo-

nent of MORB, typically comprising about 95% ofvesicle-entrapped gases. Vesicle pressures are high,notably in the popping rocks from 3000-m oceandepth that continued to explode for several days after

Žbeing brought to the surface Sarda and Graham,


.1990 . It is evident that the mid-ocean ridges are amajor source of CO recycled to the atmosphere and2

that a surge of mantle flow, such as occurred inMesozoic time, would have yielded a catastrophicincrease in atmospheric CO . Some investigators2

have attempted to identify primordial mantle carbon,based on carbon isotopic composition, but the ex-tracted CO has a wide range of d

13C and the2

evidence for a recycling model of MORB generationleads one to favor the hypothesis that the maincarbon source is recycled carbonate sediments plusminor organic carbon. The range of carbon isotoperatios can be attributed to variable C rC ratiocarb org

as well as to two fractionation processes. Outgassingof 13C-enriched CO from magma reservoirs or con-2

duits, has been proposed as a mechanism for lower-13 Žing d C from y4‰ typical of partly degassed

.ApoppingB rocks to between y6‰ and y8‰, atypical range for carbon dissolved in more thor-

Žoughly degassed MORB glass see experiments ofJavoy et al., 1978; recent reports of Blank et al.,

.1993; Pineau and Javoy, 1994 . Decarbonation reac-tions have the potential for additional fractionation inthe same direction. A dispersed carbonate melt, to beexpected where sedimentary carbonates are recycledto the MORB source, will react with mantle silicatesto generate CO at pressures less than about 18 kbar2Ž . Ž;60 km depth Hunter and McKenzie, 1989;

.Keith, 1993, p. 227 . The potential for carbon iso-tope fractionation by decarbonation reactions can bejudged from a pioneer investigation of metamorphicreactions and CO loss at the contact between an2

intrusive mass and surrounding marine limestoneŽ .Deines and Gold, 1969 . They analysed an array ofsamples extending across the contact zone andshowed a reaction-induced isotopic change from nearzero d

13C in unaltered marine limestone to y5‰ forresidual calcite within the intrusive mass.

The water content of MORB is remarkably low,Ž .;0.5 wt.% Pineau and Javoy, 1994 , consistent

with proposals that the original water content hasbeen lowered, possibly by re-equilibration in a CO -2

rich environment in upper mantle or in crustal magmaŽ .chambers Bottinga and Javoy, 1989 . A recent in-Ž .vestigation Papale, 1995 led to the conclusion that

as little as 0.1% CO in the system produces early2

exsolution of water from a melt at higher pressurethan in a CO -free system. On that basis the indi-2

cated CO content of MORB, as much as 9= the2Ž .equilibrium value Kingsley and Schilling, 1995 , is

evidently a more-than-adequate H O collector.2

2.1.3. Volatile-promoted meltingJudging from the volatile content of MORB, the

relevant phase equilibrium system for heterogeneousupper mantle melting is basalt–peridotite–CO –H O2 2Ž .Fig. 12 , not anhydrous peridotite. MORB magma isproposed to be generated by selective melting of amixture of recycled oceanic crust plus mantle peri-dotite, in which the recycled basaltic crust is thedominant component. That type of heterogeneousmelting has been proposed for the Columbia Riverbasalts, based on high-pressure melting experimentsŽ . Ž .Takahashi et al., 1998 . Kogiso et al. 1998 usedthe diamond aggregate method to collect the meltfraction, and showed that the full range of majorelement diversity of OIB, tholeiitic to alkalic, can beproduced from a heterogeneous crust:mantle mixturedominated by the crustal component.

It is proposed that the recycled LVZ mixture issubject to heterogeneous melting at a depth near 75km. High-pressure melting experiments on peridotiteŽ .Beattie, 1993 show that the observed high230 Thr238U ratios of MORB and OIB require partial

Ž .melting in the garnet stability field Fig. 12 . Therecycled masses of hydrous material will react withand hydrate the enclosing mantle and partial meltingof the mixture will yield resurgent MORB. Control-

Žling variables include depth of melting Shen and.Forsyth, 1995 , the high-pressure stability of some

Žhydrous phases Bohlen et al., 1994; Schmidt and.Poli, 1994 and variation in the H OrCO ratio2 2

ŽYoder, 1976; Eggler, 1978; papers in Yoder, 1979;.Mysen, 1987 , which affects the activity of water

and the position of the solidus. Maximum thermalstability of serpentine is about 7008 at 20 kbŽ .Iwamori, 1998 , well below the melting ranges ofperidotite and basalt. Volatile-promoted melting ofcrustqmantle mixtures can therefore be attributedmainly to dehydration of amphibole.

The proposed heterogeneous melting can be rep-resented as the melting of dolomite hornblende gar-

Ž .net peridotite DHGP of Fig. 12 . At pressures aboveŽ20–22 kb depths below the stability limit of horn-

.blende , the h –h solidus is unbuffered, in the2 4


Ž .Fig. 12. Elements of the volatile-buffered system peridotite–CO –H O after Olafsson and Eggler, 1983; Keith, 1993; Iwamori, 1998 ,2 2

combined with melting ranges of eclogite and basalt, intended to show approximate melting relations of heterogeneous mixtures of uppermantle peridotite containing blobs of metamorphosed basaltic crust. The solid line through h –h –h is a typical solidus for the hydrous3 2 4

Ž .peridotite system with hornblende present. Solidus steps are related to the falling pressure appearance of hornblende h and to2Ž . Ž .decarbonation reactions of dolomite h . At low pressures depths shallower than h , the activity of water is buffered to a low value3 2

Ž . Ž . Ža ;0.2 , by the presence of amphibole hornblende , and all of the stability fields shown in pale gray pattern are solid assemblages now. Ž .melt fraction . At pressures exceeding the amphibole-outline h –h the h –h solidus is unbuffered, and moves to lower temperatures1 2 2 4

with increasing water pressure, and melting takes place as pressures exceed the h –h line.1 2Ž . Ž . Ž .The dotted line through s and g is the dry solidus for a peridotite sample Takahashi and Kushiro, 1983 with cusps related to the

Ž .low-pressure stability limits of spinel and garnet. The example of a melting path for a rising mass of basaltic composition a–b–c–d , isŽ .modified from Yoder 1976, p. 102 . Omitted, for the sake of simplicity, are the dolomite:magnesite transition, with increasing pressure, and

Ž . 2the development of a carbonatite melt fraction Brey et al., 1983; Hess, 1993, p. 90 . The selected geotherm example, 50 mWrm , is basedŽ .on typical surface heat flow in oceanic crust, and on seismic attenuation Sato and Sacks, 1989; Jessop, 1990, p. 155 .

Ž .Dsdolomite, Gsgarnet, Hshornblende, Ls liquid melt fraction , Psperidotite, Ssspinel, Vsvolatile phase, mainly CO qH O,2 2Ž . Ž .B sbasaltq liquid, E seclogiteq liquid, expanded to lower temperatures under hydrous conditions. Pressure scale: 10 kbars1 GPa.

absence of other dehydrating minerals, and it movesŽ .to lower temperatures toward h with increased1

Žactivity of water Burnham, 1981; Keith, 1993, p..258 . The field of DGPqL is thus expanded to

lower temperatures and the melting of hornblendegarnet peridotite takes place, with temperature rise,across the Aamphibole outB line, h ™h , at a depth1 2

of about 75 km. That line also corresponds to thehigh-pressure part of the amphibole stability limit in

Ž .the basalt–H O system Liu et al., 1996 . In that2

system basalt-derived amphibolite transforms viametamorphic reactions to garnet amphibolite to am-phibole eclogite to eclogite, and the Aamphibole outBdehydration line is taken to define the melting of


Ž .hornblende eclogite Ernst, 1999 . The indicatedŽ .melting range of eclogite, E of Fig. 12 , is arbitrary,

dependent on whether the definition includes horn-blende eclogite.

ŽSome published phase diagrams Hess, 1993, p..90 show a higher Aamphibole outB pressure, near

Ž .3.2 GPa 32 kbar, Fig. 12 , but other investigatorsassociate that pressure with the stability limit of

Ž .zoisite Zack et al., 1999 . Melting experiments onŽmantle peridotite at 1 GPa Hirose and Kawamato,

.1995 showed that partial melts in systems with lessthan 2.5% water have compositions close to thosefrom anhydrous experiments at the same degree ofpartial melting. At a fixed temperature the degree ofmelting is directly related to the water content, as

Ž .shown by tests at 12508C: 6% melt dry conditions ,Ž . Ž15% melt with 0.2% H O , and 24% melt with2.0.5% H O .2

Further melting of a rising heterogeneous masscan be considered as decompression melting of basalt.Ž .Yoder, 1976, p. 101; Koga et al., 1998 . The pro-portion of liquid will increase with continued rise

Ž .through melting range B , and the magma will reachŽthe surface at near-liquidus temperatures about

.11708C .The contrast between basalt at mid-ocean ridges

versus andesite-plus-basalt at trench–arc systems canbe related to a major difference in volatile content:dominant CO at the MORB source versus dominant2

H O at the sites of melting in the Indonesian and2

circum-Pacific subduction zones. Experimental stud-Ž .ies Eggler, 1978 lead one to suggest that the calc-

Ž .alkaline andesitic trend of magmatic differentiationis inhibited by a high CO content. Most of the ridge2

crest is above carbonate compensation depth, withthe result that carbonates are the most commonsedimentary interlayers in near-axial basaltic crustŽ .Hall and Robinson, 1979 , the material that is recy-cled to the MORB source. By comparison, the rela-tively high H OrCO ratios of circum-Pacific arc2 2

Ž .volcanics Myers and Johnson, 1996 are proposed toreflect dominant hydrous minerals and a carbonatedeficiency in recycled deep-trench sediments, a re-sult of their accumulation below carbonate compen-sation depth. Additional results of the highH OrCO ratios in the arc systems include the2 2

typical occurrence of explosive volcanoes and ofhydrous granodiorite intrusive masses, some of which

are the hosts of major porphyry copper and molybde-num deposits.

The abrupt shift in petrology at ocean to continenttransition of the world rift system can be attributedto the character of the crustal material being recycledto the sub-rift region of partial melting. In oceanicregions, Gulf of Aden, for example, the dominantrecycled material is oceanic crust, whereas the sub-continent recycled material includes rift valley sedi-ments as well as residual oceanic crust recycled viathe major return-flow gyre.

2.1.4. Local sub-axial recyclingLocal recycling to the MORB source is indicated

by abrupt variations in isotopic and trace elementsignatures. Small-scale heterogeneity of the source isindicated by variable compositions of melt inclusions

Žin Mg-rich olivine phenocrysts Sobolev and Shim-.itzu, 1994 . Larger scale heterogeneity is indicated

by analysis of closely spaced samples from the EPRŽ .Prinzhofer et al., 1989; Langmuir et al., 1992 .Their results showed extreme local variations ofisotope ratios and trace element chemistry that can-not easily be attributed to a single plumbing system.Even more damaging to the conventional plate tec-tonics model are the results of sampling along aridge-normal profile. AEnrichedB chemistry, indi-cated by high BaOrTiO ratios, was found only to2

the east of the EPR axis, and there is no indication ofa symmetrical, ridge-centered geochemical pattern.

The most significant indicators of local sub-ridgerecycling are provided by correlations between re-gional variations in crustal and sedimentary materialsand their trace element and isotopic signatures inMORB. The proposed tracers of recycling are pre-sented in Fig. 13, a summary of compositional fea-tures of volcanic glass and young, least-altered basaltsfrom the axial zone of the northern MAR, the mostextensively sampled part of the mid-ocean ridgesystem. The data can be divided into three sectors inorder to simplify compositional comparisons: a

Ž .southern Sector 25–338N , taken to represent ANor-Ž .malB oceanic basalt N-MORB , an Azores Sector

Ž .35–508N that includes the Azores Platform andadjacent fracture zones, and a Reykjanes Sector that

Žencompasses Reykjanes Ridge and Iceland 50–.708N . The Azores platform and Iceland, culmina-

tions of those sectors, are elevated regions of the


Fig. 13. Geochemical profiles of basalt and basaltic glass from the axial zone of the Mid-Atlantic Ridge, 25–758N, including the IcelandicŽ . Ž . Ž) .Ic and Azores Az anomalies and compositional spikes related to nodal deeps at ridge–transform intersections of major fracture

Ž . Ž . Ž . Ž . Ž . Ž .zones: Jan Mayen JM , Charlie–Gibbs G , Kurchatov Kv , Oceanographer O , Hayes H and Atlantis A , and an unlabelled FractureŽ . Ž .Zone at 458N. The profiles are redrawn from Schilling 1986 and Melson and O’Hearn 1986 . Depths are reported sampling depths.

LarSm data are normalized to chondrites, helium isotope ratios to the atmospheric ratio. Dashed lines typical helium isotope ratio ofN-MORB. FeO)s total iron recalculated as FeO.

ridge system that are characterized by relatively highrates of volcanism, evidenced by thicker than aver-

Žage basaltic crust Melson and O’Hearn, 1986, p..118 and by quite different geochemical anomalies

Ž .relative to AnormalB MORB . The high rates ofvolcanism have previously been attributed to mantleplumes and the contrasted geochemical anomalies

either to separate undepleted deep sources of theŽ .plumes Schilling, 1986, p. 144 or to plume-derived

magmas that were fractionated during long residenceŽtime in crustal magma chambers Melson and

.O’Hearn, 1986, pp. 129, 134 . It was implied, in theabove reports, that there must be a fundamentalcompositional difference at the source. That implica-


tion is consistent with the conclusion of SigurdssonŽ .1981 . He showed, in a study of major elementvariation in North Atlantic basaltic glass samples,that they can be divided into two groups, roughlyequivalent to the Reykjanes and Azores sectors, andthat major element variations within each group de-fine a fractionation trend from olivine tholeiite toquartz tholeiite. Sigurdsson showed that the groupfractionation paths are clearly separated and that

Žsome of the major element contrasts iron content,.for example persist from primitive to evolved basalts

and require compositional differences at the source.One can appeal to mantle heterogeneity, for which

Žthere is considerable evidence Melson and O’Hearn,.1986, p. 130 , but it is difficult to account for the

observed gradual changes in geochemical signaturesby a process so inherently slow and inefficient as themixing of viscous mantle plumes. It is proposed, and

Ž .supported by geochemical evidence Fig. 13 that theindicated heterogeneous source of MORB is theresult of local recycling of basaltic crust plus vari-able proportions of interlayered sedimentary compo-

Ž .nents Kurz et al., 1982 .In the proposed recycling model of oceanic vol-

canism, there is a correlation between the rate ofvolcanism and the compositional features of MORB.The voluminous volcanism of the Azores and Ice-landic regions is attributed to relatively rapid convec-tive turnover and to high rates of recycling of low-

Ž .melting crustal materials cf. Schilling, 1986 . Theunderlying causes of increased recycling are judgedto be different, however. The Azores platform islocated at a triple junction, conventionally defined asthe intersection of the Mid-Atlantic Aaxis of spread-ingB with a zone of dextral shear between the

ŽEurasian and African plates Vogt and Tucholke,.1986, pp. 370, 418 . In the proposed viscous flow

model, a triple junction, such as that within theAzores platform, is the conjuncture of three convec-tion cells, a location where cell-boundaries cometogether and yield a focused local zone of strongdownwelling, recycling and volcanic activity. Thenorthward shallowing of Reykjanes Ridge is alsoattributed to an increase in rate of volcanism and toaccentuated uplift due to increased volume of low-density crust–mantle mixture below the ridge. In theReykjanes Ridge case, a proposed northward in-crease in rate of downwelling is attributed to pro-

gressive northward narrowing of convection cellwidth, as indicated by the distance from ridge axis todominant upwelling zones beneath adjacent conti-nental shields. Additional focusing of strong down-flow can be attributed to prolonged cumulative buildup of a cold residuum below the most northerly

Ž .MAR cf. Fig. 4 , preserved because it was relativelyprotected from disruption during the Pacific-centered

Ž .Mesozoic mantle surge Keith, 1993 .

2.1.5. Sites of sediment accumulation and recyclingClues to interpretation of the North Atlantic geo-

Ž .chemical anomalies Fig. 13 are provided by severalclasses of covariance. The covariant LarSm and Srisotope anomalies are attributed to local increases inthe recycling of basaltic crust with an above averageproportion of interlayered terrigenous sediment, rela-tively enriched in large-ion elements and radiogenicisotopes. The observed covariance of those anoma-lies with present topography and crustal thicknesscan be attributed to a common dependence of vol-canic rate and composition on regional increases in

Ž .the rate of recycling see below: anomaly spikes . Asecond class of covariances includes those relatinggeochemical anomalies of North Atlantic basalt withtheir proximity to RFZ intersections, equated withthe confluence of roll boundaries and the majorsub-axial convection cell boundary. Some investiga-tors have referred variations of ridge features to RTI,but RFZ intersections are emphasized here becausesome of the significant anomalies are associated withnon-transform fracture zone intersections, i.e. wherethere is no apparent ridge offset. The RFZ associa-

Žtion is most clearly defined in the TiO profile Fig.2.13 but is evident in the other selected geochemical

Žprofiles, except that for helium isotope ratio see.below . The observed associations are attributed to

the accumulation of sediments in fracture zone val-leys and in nodal deeps at RFZ intersections, and tostrongly focused recycling of basalt and interlayeredsediment at those nodes. The FeO and TiO spikes,2

for example, are attributed to the accumulation atRFZ nodal deeps, of fine-grained chemically resis-tant titaniferous magnesiochromite and magnetite

Ž .fractions Roeder, 1994, p. 734 derived from weath-ering of basalt during the sub-aerial stage of theformerly elevated MAR. It will be evident, in accordwith the proposed recycling model, that the size and


definition of RFZ spikes will depend on the exis-tence of a fracture zone valley and nodal deep, bothof which have variable expressions. Prime examples,the well-defined Jan Mayen and Charlie–Gibbs geo-chemical spikes, are associated with deep fracturezone valleys, 1000–3000 m below adjacent trans-

Žform and axial Saemundsson, in Vogt and Tu-.cholke, 1986 , more than 2000 m deeper, for exam-

ple, at the eastern end of the Kane transformŽ .Auzende et al., 1994 .

2.1.6. Sedimentary source of oceanic sulfide depositPb

The case for local recycling is strongly broughtout by a comparison of distribution and isotopiccomposition of lead in basalt, black smokers, oceanic

Žsulfide deposits and local sediments LeHuray et al.,.1988; Doe, 1994 . Lead is notably rare in black

smokers and sulfide deposits on sediment-starvedridges, and the traces of lead have isotopic composi-tions matching that of the local basalt. In contrastwith a sediment-starved ridge, lead is consistentlypresent where there is an abundant supply of terrige-nous sediment and at those sites the Pb isotopic

composition of the galena matches that of the ter-rigenous clastic sediment. A prime example is pro-vided by data for the lead isotopic composition ofsulfide deposits on sediment-starved Juan de FucaRidge and sediment-swamped Gorda Ridge. The lat-ter shows a close match with the isotopic composi-

Žtion of lead in the local terrigenous sediment Fig..14 . In addition, the Pb isotope trend-line of Gorda

Ridge basalt indicates that the local sediment is asignificant component in the source of basalticmagma.

In summary, it is proposed that the geochemicalanomalies of MORB can be attributed to local recy-cling of altered basaltic crust plus sedimentary inter-layers characterized by regional variability in the

Ž .proportions of the following components: 1 terrige-Ž .nous sediment, 2 FeqTi oxide minerals residual

from chemical weathering of exposed basaltic ter-Ž .rains, 3 soils and near-shore precipitates enriched

Ž .in Fe, also from chemical weathering of basalt, 4weathered rubble of volcanic rocks from formerlyexposed mid-ocean terrains. If the recycling modelfor MORB genesis is essentially correct, it wouldseem to call for a review of global geochemical

Fig. 14. Partial record of lead isotopic composition of basalt, sulfide deposits and sediments from two dredge sites on sediment-swampedŽ .Gorda Ridge eastern Pacific compared with the Pb isotopic range for basalts from sediment-starved Juan de Fuca Ridge. Redrawn from

Ž .Figs. 2 and 3 of LeHuray et al. 1988 .


budgets, e.g. the CO budget, and the sedimentary2

source of oceanic ore-forming fluids.

2.2. Global geochemical systems

2.2.1. Steady-state and episodic recyclingThe essence of the proposed global model is that

oceanic crust is a principal reservoir and that selec-tive recycling of its components is a counter toweathering and riverine transport from continent toocean, and a key process in the self-regulating Earthsystem. The sediment-recycling aspect of the pro-

Žposed model is similar to that of Armstrong 1968,.1991 , but modified to emphasize that the mid-ocean

ridge axis is the main site of persistent recycling, onethat has been supplemented, in post-Cretaceous times,by subduction at the transient Indonesian and cir-cum-Pacific arc systems. Recognition of the long-term dominance of mid-ocean recycling of oceaniccrust and interlayered sediment allows dismissal oftwo principal objections that induced prior authors tominimize the importance of sediment recycling: thenon sub-ductibility of low-density sediments and thelimited subduction at Indonesian and circum-Pacificarc systems. Sialic melt fractions, derived from sedi-ments, will not normally be erupted or intruded atthe ridge axis because their penetration of water-cooled oceanic crust is inhibited by their high viscos-ity. Granitic melts are estimated to be about 1000

Ž .times stiffer than basaltic melts Elder, 1976 . As aresult, the sialic magmas will tend to be incorporatedin return flow gyres and eventually in upwellingplumes beneath continental shields and borderingmobile zones. Crustal recycling and consequent de-velopment of a Amarble-cakeB mantle have beenproposed by several investigators, including Lewis

Ž .and Smith 1997 , who based their recycling modelon the Nd–Hf isotopic array of MORB. It is gener-ally assumed that there is a long-term balance be-tween erosion and continental regeneration by recy-cling of sediments. The effects of recycling on globalgeochemical systems depend on the selectivity of theprocess and on differences of selectivity duringsteady-state flow and episodic surges of mantle flow.During surges, large masses of oceanic crust willremain below melting range in the sub-axial subduc-tion zone and eventually will be recycled to sub-con-tinental regions, thus providing a source for large

Žeruptions of continental flood basalt cf. Takahashi et.al., 1998 .

In accord with the proposed model, the averagetrace element and isotopic signatures of present-dayMORB are taken to be the integrated result of long-term recycling, and departures from that average willbe evident only where the local recycling mix issignificantly different. Examples are provided by

Ž .isotopic anomalies in North Atlantic basalt Fig. 13 ,Ž .and in Gorda ridge basalt Fig. 14 whose Pb iso-

topic trend is attributed to recycling of young terrige-nous sediment, relatively enriched in radiogenic lead.

Major surges of mantle flow, such as the one thatapparently began in early Mesozoic and peaked inCretaceous time, involve increased rates of recy-cling, partial melting and volcanism. Increased vol-canic activity presumably supplied greatly increasedvolumes of volcanic gases to the atmosphere. Anincreased flux of chlorine and fluorine is of specialinterest because of their known effects in destroyingthe ozone layer. Perhaps it is time to reconsider thehypothesis that repeated volcanic destruction of theozone layer affected the selectivity of Cretaceous

Ž .extinctions Keith, 1982 .

2.2.2. The carbon budgetThe proposed recycling model leads to the corol-

lary that mid-ocean recycling probably provides amajor source of variable carbon dioxide additions tothe atmosphere. The ocean ridge sites, including rollmargins and the axial zone, are normally at depthsabove calcite compensation, with the result that car-bonate sediments are the dominant sedimentary ma-terial being recycled. That condition is in contrast tothe recycling and partial melting regime at inclinedsubduction zones of Indonesia and the Pacific rim,where the sediments available for recycling are accu-mulated in trenches at 9 to 11 km depths, well below

Ž .calcite compensation depth ;4.2 km , and there-Ž .fore tend to be carbonate deficient cf. Section 2.1.3 .

The final stage of a mantle surge, the beginningof a return to steady-state conditions, will involve agradual reduction in the volume of sub-ridge subduc-tion mixture, as the system returns toward a balancebetween mid-ocean subduction and ocean to conti-nent return flow. Gradual isostatic subsidence ofmid-ocean ridges will be accompanied by continentaluplift, a result of sialic additions to the continents.


The combined effect will show up as a lowering ofsea level and an increase in land area, an example ofwhich is provided by the abrupt increase that peakedat the end of Cretaceous time, and continued gradu-ally through Tertiary, to a total of about 40%. Hereagain, there will be a major effect on the atmosphericCO budget. The system clearly is complex and2

Žseveral proposed models Bickle, 1996; Berner andCaldeira, 1998; Kerrick and Caldera, 1998, papers in

.Ruddiman, 1997 , have placed variable emphasis oncarbon dioxide sources related to mantle and meta-morphic degassing plus oxidation of organic matter,and on carbon dioxide removal by acceleratedweathering of silicates during global greenhousewarming. Increased weathering of newly exposedand uplifted land areas presumably will cause amajor reduction in atmospheric CO . The suggested2

modification to previously proposed models is thatocean ridge processes probably should be consideredas one of the dominant controls on the atmosphericCO budget. In view of the carbonate deficiency in2

sediments of the trench–arc systems, it probably isincorrect to assume that CO contributions from2

volcanic arcs are as large as those from mid-oceanŽ .ridges Marty and Tolstikhin, 1998 .

2.2.3. The lead paradoxThere is a long-recognized discrepancy between

Ž .the measured ThrU ratio of MORB about 2.5 andthe ThrU ratio calculated from the 208Pbr204 Pb and206 204 Ž . ŽPbr Pb isotope ratios about 3.7 Zartman and

.Haines, 1988 . Lead in volcanic rocks, presumed bymany investigators to represent mantle lead, is appre-ciably more radiogenic than bulk Earth values, im-

Ž238 206 .plying that mantle m Ur Pb has increasedover time. The Alead paradoxB evidently requires aU-enriched reservoir for Pb isotopes. The paradox isemphasized by an opposite paradox regarding Srisotopes in volcanic rocks; 87Srr86 ratios are lowerthan expected and apparently require an Rb-depletedreservoir. Both paradoxes can be accounted for bysediment recycling in which the U-enriched reser-voirs are continental crust, terrigenous sediment andblack shales, whereas the principal Rb-depletedreservoir is limestone. The proposed model is similarto the recycling simulation of Armstrong and HeimŽ .1973 , devised to account for evolution of the Pband Sr isotopic systems, and to the models of Chase

Ž . Ž .1981 and Zindler et al. 1994 based on differentialŽ .recycling of sediments. White et al. 1984 suggested

that radiogenic Pb may be recycled by sedimentŽ .subduction, and Zartman and Haines 1988 pro-

posed preferential recycling of U relative to Th, aresult of the mobility of uranium and the insolubilityof thorium. Data for m values of marine sediments

Žinclude measurements on terrigenous sediments m

. Ž .s8.7 , pelagic red clay and biogenic ooze ms4Ž . Žand manganese nodules ms0.7 White et al.,

. 6q1984 . Uranium is oxidized to U during weather-ing and is soluble in seawater as a carbonate com-plex. It is collected on the organic carbon com-pounds of marine sediments, notably in black shales,taken to be the largest sink in the global budget of

Ž .uranium Klinkhammer and Palmer, 1991 . Uraniumis also added to basaltic crust by hydrothermal circu-lation of seawater. The black shales and alteredbasaltic crust are judged to be the principal high-mmaterials that are recycled to the magma source andyield the enhanced uranium content and anomalouslylow ThrU ratios of MORB. Extreme uranium en-richment, which produces the HIMU class of oceanicisland basalts, may involve the recycling of organic-rich black sapropels to the magma source. The HIMUsediments presumably are interlayered in volcaniccrust formed in Cretaceous times, when such sedi-ments were abundantly deposited in stagnant oceanbasins.

2.2.4. HeaÕy noble gases in oceanic basaltThe noble gas geochemistry of oceanic basalt is

Žsummarized in recent papers Honda et al., 1991;.Valbracht et al., 1997; Moreira et al., 1998 , along

with citations to the many earlier investigations ofŽMORB glass Jambon et al., 1985; Marty and Oz-

. Žima, 1986; Fisher, 1994a,b , and of OIB glass Al-.legre et al., 1983; Kaneoka et al., 1986 . Special

emphasis has been accorded to Loihi seamount,Hawaii, where the lavas have not been exposed at

Žthe surface Sarda et al., 1983; Staudacher et al.,.1986 . Principal features of the noble gas inventory

in basaltic glass samples include atmospheric ratios38 36 Ž .of Xe, Kr and Arr Ar Hiyagon et al., 1992 ,

together with a wide range of 40Arr36Ar ratios, fromŽ .near-atmospheric 295 to values over 25,000.

Despite the indications of air-like heavy nobleŽgases, several investigators e.g. see Staudacher et


.al., 1986 , have attributed the noble gas inventory toa primordial source in the lower mantle, presumed tobe only slightly outgassed and to have retained noblegas ratios similar to those of air. Patterson et al.Ž .1990 discounted the hypothesis of a lower mantlesource and proposed that atmospheric contaminationof sub-oceanic basalts occurs via noble gases dis-solved in seawater. The seawater path offers onemechanism for atmospheric contamination but doesnot account for the wide range of argon isotoperatios, atmospheric to radiogenic. I propose an alter-native pathway, the recycling of terrigenous clasticsediments that are intercalated with oceanic basalt.Proposed argon sources include detrital, weather-re-sistant igneous and metamorphic minerals such asfeldspars and micas, together with the clay mineralsof soil profiles, including kaolinite, smectites andillites, that presumably collect atmospheric argon by

Ž .adsorption and end up in shales Dong et al., 1995 .Incremental additions of atmospheric noble gaseshave been shown to accompany the weathering of

Ž .chondritic meteorites Scherer et al., 1994 , and Hur-ley showed, more than 30 years ago, that clay miner-als retain radiogenic argon and thus provide for

Ž .K–Ar dating of sediments Hurley, 1966 . Addi-tional sources of radiogenic argon include authigenic

Žpotassic minerals glauconites in shales, K-feldspar.in calcareous rocks , minerals that generate radio-

genic argon within the sedimentary environmentŽ .Steinitz et al., 1995 . Recent tests of the release ofargon by stepwise vacuum heating of illite and glau-

Ž .conitic clay Hassanipak and Wampler, 1996 showedthat radiogenic argon is in structural sites in clayminerals and is more tightly held in potassic illitesthan in iron-rich glauconites. It is worth noting thatpotassic illites transform to more stable mica struc-tures on metamorphism. It seems reasonable to ex-pect that the argon will be retained in continent-de-rived detrital and hemipelagic sediments that arerecycled as interlayers in oceanic crust. There is anapparent need for analyses of other noble gases inclay minerals of soil profiles.

The sediment recycling model has the advantagethat it can account for the coexistence, in oceanic

Ž .basalt, of atmospheric argon mainly radiogenic ,and variable proportions of new radiogenic argonfrom continental sediments. The proposed modelgains support from the observation that MORB from

the MAR and EPR have wide ranges of argon iso-topic compositions, consistent with the general avail-ability of terrigenous sediments in the recycling mix.In contrast, mid-Pacific basalts, exemplified by thesamples from Loihi seamount, are in a region ofsparse and invariant terrigenous sediment, and theresultant argon isotope ratios are in a low narrowrange. Typical ranges of 40Arr36Ar ratios, from FisherŽ . Ž .1986, 1989 , Hiyagon et al. 1992 , Marty and Hum-

Ž . Ž .bert 1997 , Sarda et al. 1985 , and Valbracht et al.Ž .1997 are:

Mid-Atlantic Ridge, 11y408N:

371y42,360 Ns52 medians2417Ž .

East Pacific Rise and Galapagos:

308y17,800 Ns63 medians2660Ž .

Mid-Pacific and Hawaiian Islands:

316y2701 Ns29 medians794Ž .

The helium and neon inventories of oceanic basaltŽ .are considered separately Sections 2.4.3 and 2.4.4 ,

on the grounds that helium is lost from the atmo-sphere and also, with neon, is more mobile than theheavy noble gases, thus subject to homogenizationby diffusion. The proposed recycling of heavy noblegases is based on two assumptions: firstly, that conti-nental sediments are the principal source, and sec-ondly, that partial melting and the generation ofbasalt constitute an efficient process for collecting

Žnoble gases. A recent review McDougall and Honda,.1998, p. 180 assumed that the solubilities of noble

gases in silicate melts are extemely low, but experi-mental determinations of partition coefficients be-tween olivine and melt show the opposite, that thenoble gases are efficiently collected by a partial melt

Žfraction Hiyagon and Ozima, 1986; Valbracht et al.,.1994 . Distribution coefficients for olivine:basaltic

melt pairs are K F0.07, K s0.05–0.15. Confir-He Ar

mation of the early introduction of atmospheric ar-gon to a magma source was provided by Farley and

Ž .Craig 1994 . They analyzed argon concentrationsand 40Arr36Ar ratios in olivine phenocrysts from anOIB and showed that contamination of the magmawith atmospheric argon must have occurred prior toolivine crystallization.


2.3. ReserÕoirs and recycling

2.3.1. Geochemical constraints on the deep-plumemodel

The deep-source plume model of oceanic islandvolcanism is not only faced with contrary geophysi-

Ž .cal evidence Section 7.01 , it is also inconsistentwith several critical geochemical features of OIB:

Ž .1 widely different isotopic signatures of adjacentoceanic islands, for example those of Mangaia, Atiu

Žand Aitutaki in the Cook–Austral chain Palacz and.Saunders, 1986 ;

Ž .2 significant isotopic contrasts between volca-noes, even within one island, for example betweenthe Loa and Kea linear volcanic chains, 40 km apart

Ž . Ž .on Hawaii Kurz et al., 1995 Fig. 15 ;Ž .3 along-chain isotopic contrasts, e.g. between

Mauna Loa and the extinct volcano Kahoolawe,Ž .about 100 km to the northwest Fig. 15 ;

Ž .4 rapid temporal changes of isotope ratios, e.g.in the Sr and Nd ratios of lavas from the Mauna Loa

Ž .borehole Rhodes and Hart, 1995 ;Ž .5 variable disequilibria and implied small-scale

heterogeneity in the U and Th systems within theŽlava sequence of a single volcano Cohen et al.,

.1996 ;Ž .6 variable helium isotope ratios in an age se-

Ž .quence of Hawaiian volcanoes Fig. 19 and system-atic changes within the lava sequence of a single

Ž .volcano Fig. 20 .

2.3.2. Principal reserÕoirsThe proposed recycling model for oceanic basalt

Ž .is based on an array of isotopic indicators Fig. 16 ,together with evidence that oceanic crust is not lostby subduction at Pacific and Indonesian subduction

Ž .zones plate tectonics model , but rather is preservedŽ .by load-induced mid-oceanic subsidence Section 1

Ž .Fig. 15. Outline map of Maui and Hawaii after Rison and Craig, 1983 , to show two separate lines of volcanoes, commonly named theAKeaB and ALoaB trends. Within each linear array, volcano ages decrease toward the southeast.


Ž .Fig. 16. Strontium and lead 206r204 isotopic composition of oceanic island basalt and MORB to show the divergence of isotopic trendsŽ .from a central focal composition F equated with a bulk long-term reservoir, an accumulated crustrmantle mixture in sub-ridge upper

Ž . Ž . Ž . Ž .mantle see text: FOZO and seismic LVZ . The figure is redrawn from Davies et al. 1989 , Keith 1993 and Hofmann 1997 withŽ . Ž . Ž . Ž . Ž . Ž .additional details from Chaffey et al. 1989 St. Helena , Chauvel et al. 1992 Tubuai , Farley et al. 1992 Tutuila ; Palacz and Saunders

Ž . Ž . Ž . Ž . Ž . Ž . Ž . Ž . Ž1986 Rurutu ; Staudigel et al. 1984 and Stille et al. 1986 Hawaii ; Vidal et al. 1984 Marquesas ; Weis and Frey 1991 Ninetyeast. Ž . Ž .Ridge ; Woodhead and Devey 1993 Pitcairn seamounts .

Ž .The broad trend of Pacific MORB is attributed mainly to mixture of F with depleted lower mantle, DM see text . Other compositions anddivergent trends of oceanic basalt are attributed to mixtures of locally recycled OIBqMORB, with one or more of the followingsedimentary components: TCs terrigenous clastic sediments; PELAGICSspelagic sediments dominated by marine precipitates; MPs

Ž .Mesozoic pelagic sediments, mainly carbonates see text ; HIMUsSubduction mixtures dominated by black sapropel with abundantorganic carbon compounds and enhanced uranium content.

Ž .Solid circles within the field of Pacific MORB are isotopic compositions of Cretaceous crust from ODP Site 843 King et al., 1993 . Loa,Kea and Koo within the isotopic field of Hawaiian lavas represent characteristic isotopic compositions exemplified by Mauna Loa, Mauna

Ž .Kea, and Koolau. Previously proposed end-member reservoirs Zindler and Hart, 1986 include: EM1 and EM2senriched mantle, DMŽ . Ž .sDMM sdepleted mantle source of MORB, HIMUsmantle source with high m high UrPb ratio .

and by eventual incorporation of lower crustal layersinto an upper mantle mixing zone, by recycling atthe boundaries of convective rolls in the upper man-tle. It is that mixture, equated with the seismicsub-ridge LVZ, that is the principal source of newbasaltic magma. Dominant components of the LVZ

Ž .mixture are oceanic crust qsediments and upper

Žmantle. In addition, there are indications Section.2.3.3 of a Mesozoic episode in which the crustal

and upper mantle components were mixed with lowermantle.

Sediment recycling to the source of oceanic basalthas been proposed to account for regional contrasts

Ž .in basalt chemistry Dupre and Allegre, 1983 and


for various features of the isotopic and trace elementŽarray Chase, 1981; Hofmann and White, 1982; Zart-

.man and Haines, 1988; Fisher, 1989; Weaver, 1991 .Ž .Zindler and Hart 1986 proposed that the isotopic

variations of OIB can be attributed to contributionsfrom four principal end members: DMMsdepletedmantle, EM1 and EM2senriched mantle reservoirs,and HIMU, a proposed mantle reservoir character-

Ž .ized by high m values high UrPb ratios . Hart et al.Ž .1992 followed up by proposing a focal zone com-

Ž .ponent FOZO , located centrally within the isotopicarray defined by the four original components. Theysuggested that the FOZO reservoir may be in thelower mantle or in a core:mantle boundary layer. An

Ž .alternative central component PHEM , defined asŽ .primitive helium mantle Farley et al., 1992 , has a

central composition similar to that of FOZO. WeaverŽ .1991 suggested that the EM1 and EM2 end mem-bers can be attributed to recycling of oceanic crust

Ž .containing small amounts of pelagic sediment EM1Ž .or terrigenous sediment EM2 . All of the above

models postulate that the sources of oceanic basaltare isolated mantle reservoirs in which the recycledmaterials accumulate and maintain their integrity forlong periods of time. Some models include a shallow

Ž .reservoir for MORB and a deep corermantleboundary reservoir for OIB.

In contrast to the conventional mantle reservoirmodel, the proposed model for OIB postulates thatthe controlling process is local shallow recycling ofoceanic crust and contained sedimentary layers, ratherthan previously proposed remote subduction attrench–arc systems followed by long-term isolationof end-member components in separate reservoirs.The proposed mechanism is load-induced subsidenceof each volcanic island system into an underlyingheterogeneous zone of volatile-induced melting. Theobserved isotopic and trace element diversity is at-tributed to regional variations in the character ofentrained sediment and to variable contributions fromthe mantle and from upper and lower layers of thesubsiding volcanic pile. I have chosen to representthe isotopic diversity of OIB by Sr and Pb isotope

Ž .ratios Fig. 16 because they exhibit a variety ofcorrelations, both positive and negative, and theirgeometric relations provide more useful diagnosticinformation than systems in which the correlations

Ž .are consistently negative Sr versus Nd , or consis-

Ž206 204 207 204 .tently positive Pbr Pb versus Pbr Pb .ŽThe NdrPb isotopic diagram Staudigel et al., 1991;

.Janney and Castillo, 1999 is as usefully diagnosticas the SrrPb array, and is free of the seawateralteration that distorts the Sr isotopic signature.

A distinctive feature of the Pb–Sr and Pb–Ndisotopic arrays is that most of the isotopic trend linesdiverge from a centrally located focal componentŽ .F , analogous to the FOZO component of Hart et al.Ž .1992 , but adapted to the current model by discard-ing the deep-reservoir concept and emphasizing evi-dence that F has the bulk composition of the sub-ridge crustrmantle mixing zone, the time-integratedresult of long-term shallow recycling. The principalF reservoir is equated with the seismic low-velocityzone but may include metamorphosed lower crust,not yet incorporated into the sub-ridge mixture. Iden-tification of the F reservoir as an ancient but stillactive mixing zone, strongly influenced by recycledsediments, is supported by the observation that its Pbisotopic composition, defined by intersection of lin-ear oceanic island trends, matches two independentestimates of average Pb isotopic composition ofoceanic sediments, including an average of pre-Holo-

Žcene sediments from the North Atlantic Hamelin et.al., 1990 , selected to be relatively free of anthro-

pogenic lead pollution, and an average compositionof sediments being subducted around the PacificŽ .Plank and Langmuir, 1998 . The following tabula-tion gives the approximate isotopic composition ofthe F reservoir, based on trend intersections and

Žpublished regression lines Davies et al., 1989;Staudigel et al., 1991; Vidal, 1992; Hauri et al.,

.1996, 1997; Hofmann, 1997 :87Srr 86 Srs0.703 206 Pbr 204 Pbs19.2143Ndr 144 Nds0.5129 207Pbr 204 Pbs15.55

187Osr 188Oss0.124 208Pbr 204 Pbs38.6

2.3.3. Episodic lower mantle contributionPacific MORB defines a broad trend from F

Ž .toward a component DM , characterized by less-ra-diogenic Sr and Pb. DM is taken to be depletedmantle, as originally proposed by Zindler and HartŽ .1986 . Subsequent investigators usually have as-sumed it to be uppermost suboceanic mantle, but astrong possibility, based on the evidence for a Pa-


cific-centered Mesozoic surge of mantle flow, pre-sumably triggered by a cluster of meteorite impacts

Ž .in the western equatorial Pacific Fig. 22 , is that theDM reservoir is reactivated lower mantle, availableduring a surge-promoted change from layered to

Ž .whole-mantle convection Section 2.5 . Evidence thatŽ .DM and DMM probably represent lower mantle is

provided by the isotopic signatures of the oldesttholeiitic basalts, Jurassic basalts from ODP Site 801in the Pigafetta Basin, near the western end of the

Ž .residual lower mantle plume Castillo et al., 1992 .Those basalts, with maximal radiometric ages of 167Ma, have extremely low Sr isotope ratios, below0.7025, taken to indicate that they came from amantle source with long time-integrated depletion of

Ž .Rb relative to Sr Castillo et al., 1992, p. 409 . Alower mantle source is favored by the obvious con-straint that an extreme Rb depletion cannot persist inthe upper mantle because it is subject to multiple-stage recycling of Rb-rich terrigenous detrital sedi-ments. It has also been suggested, based on isotoperatio signatures, that the Dupal anomaly of the south-ern ocean basalts has a lower mantle componentŽ .Castillo, 1988 .

Deep roll-margin recycling associated the pro-posed impact-triggered change from layered to wholemantle convection, and resultant mantle surge isdeduced to have formed a range of mixtures of Fwith DM, a mixture that apparently constituted asource of Cretaceous basaltic magma and oceaniccrust. Examples are provided by Cretaceous crustintersected by drilling at Site 136, to the west of

Ž .Hawaii King et al., 1993 , and by early Mid-AtlanticŽ .MORB Janney and Castillo, 2000 . Isotopic compo-

sitions of Pacific Cretaceous are shown as filledcircles in Fig. 16, and similar crust is presumed tooccur at depth below the Hawaiian islands. Theproposed model involves persistence of Mesozoiclavas in the lower crust, and their recycling into themantle at roll margins, with consequent accumula-tion of a large reservoir of old crust in a sub-ridgezone of mixing and hydration. The mixture is pro-posed to be a major component of newly formingMORB and OIB magma. Principal isotopic signa-

Ž .tures include low Sr isotope ratios see below , andhigh 3Her4 He ratios, attributed to the abundant sub-aerial basalt of Mesozoic times. Indications that un-derlying Mesozoic crust is a component in genera-

tion of some OIB are provided by the observationthat strontium isotope ratios of Hawaiian tholei-ites are in the range 0.7040–0.7045, but those ofdeeper-source alkalic basalts are in the range0.7030–0.7035. That contrast can be attributed tocontributions from DM and from Mesozoic marine

Žcarbonates intercalated with lavas of that time Sec-.tions 2.3.3 and 2.3.7 .

2.3.4. Sediment signaturesThe geochemical signatures of marine sediments

reflect mixtures of continent-derived clastics, organicmaterials and marine precipitates, modified by ad-

Žsorption and various diagenetic reactions cf. Fagel.et al., 1997 . In order to simplify a review of sedi-

ment recycling to the magma source and of theresultant isotopic and trace element signatures inoceanic basalt, it is assumed that AterrigenousB im-plies mainly clastic sediments, of all grain sizes, andthat ApelagicB implies dominant marine precipitates,including carbonates, opal, phosphates, FeMn oxy-hydroxides, etc.

Pacific MORB data, and some of the AtlanticMORB data fall along the F to DM trend, taken toimply that the dominant control is the intermixture of

Ž .average oceanic crust F crust with depleted mantle.The Atlantic data extend over a wide range, how-

Žever, and the Indian data even wider Hofmann,.1997 , presumably reflecting variable large contribu-

tions from terrigenous sediments. Much of the cen-tral Pacific basin receives limited amounts of terrige-nous clastics, due to blocking by island arcs andback basins, and to the great distances of mid-Pacificislands from continental sediment sources. As a re-sult Pacific OIB shows a dominant effect from recy-cling of marine precipitates. The proposed generationof MORB and OIB by a process involving localrecycling of crustal materials is supported by iso-topic and trace element signatures of several differ-ent sedimentary components, including marine pelag-ics but dominated by terrigenous clastic sediments,TC, whose isotopic signatures are presumed to re-flect time-integrated parent:daughter ratios.

The MP component, located approximately by theintersection of isotopic trend lines of individual is-

Ž .lands Gough, Walvis, Pitcairn and by a linearŽisotopic trend of Hawaiian picrites Bennett et al.,

.1996 , is tentatively identified as a mixture of DMq


Mesozoic pelagics, dominated by biogenic carbon-ates. Carbonates deposited from modern seawaterhave Sr isotope ratios near 0.709, too high for theMP component, but the Sr isotopic composition ofseawater was as low as 0.7068 in Jurassic time and

Ž .Veizer 1989 showed that the diagenetic changeŽfrom carbonate ooze aragonite and magnesian cal-

. Ž .cite to chalk low-Mg calcite involves a furtherlowering of the Sr isotope ratio by about 0.001,sufficient to account for the MP ratio, slightly below0.706. The low Pb isotopic composition of MP,essentially that of DM, is consistent with the pro-posal that the carbonate component dominates the Srfingerprint but acts only as a diluent to the Pbcontent of the mixture, with the result that the Pbisotope ratio is defined by that of DM. A reasonablespeculation is that the mixture formed as a result ofthe major Mesozoic disruption of Pacific mantleconvection, possibly by a cluster of meteorite im-pacts that triggered a change from layered to whole-

Ž .mantle convection Section 2.5 . If the MP compo-nent is adopted, EM1 may not be particularly useful.EM2, possibly also redundant, seems to be merelythe average composition of recycled sediment.

The divergent isotopic trends of oceanic islandŽ .basalt, mainly radiating from F Fig. 16 are at-

tributed to local subsidence of the island volcanoes,in response to volcanic loading, and to eventualincorporation of the subsiding crustal pile into asub-island zone of volatile-promoted partial melting.Local rather than remote recycling is favored by thecharacteristic isotopic differences between islands

Ž .within one group, for example Fig. 16 between UaŽPou and Nuku Hiva in the Marquesas Vidal et al.,

.1984 , between Sao Miguel and Terceira in theAzores, and between the southern and northern is-land groups of the Cape Verdes Archipelago.

The HIMU component is proposed to comprisemixtures of F with stagnant basin sapropel contain-ing abundant organic carbon compounds, relatively

Ž .enriched in uranium Barnes and Cochran, 1988 ,Žand with consequent high m values high UrPb

.ratios . It is equivalent to the HIMU component ofŽ .Zindler and Hart 1986 but without the implication

of an isolated deep-mantle reservoir. There are indi-cations that organic-rich black muds were abun-

Ždantly formed in Cretaceous stagnant basins Keith,.1982 , but sediment interlayers of that character are

not restricted to Mesozoic crust. The compositionalfeatures assumed for the postulated HIMU sediment:lack of bioturbation, high iron content, sulfates withlow d

34S, are similar to those of black sapropels ofthe Kupferschiefer, the Upper Permian Zechstein

Žsequence Margaritz and Turner, 1982; Sun et al.,.1995 , and the Permian Phosphoria formation of the

Ž .northwestern United States Piper et al., 1994 . Simi-lar black sapropels of the Mediterranean are at-

Ž .tributed to mat-forming diatoms Kemp et al., 1999 .The previously puzzling high CaO content of HIMU

Ž .basalt Kogiso et al., 1998 probably can be at-tributed to gypsum-bearing sediment. The PacificHIMU islands, such as Rurutu, Tubuai and Mangaia,

3 4 Žhave extremely low Her He ratios Hanyu and.Kaneoka, 1997 , presumably due to an excess of

uranium-derived helium-4.A possible common denominator is that the prin-

cipal HIMU islands are adjacent to formerly activefracture zones, sites of roll margin recycling, whoseactive stage typically is marked by deep valleys,preferred collection sites for deep stagnant-basin de-posits. The HIMU islands of the Pacific Australchain are located adjacent to the Challenger FZ, along, formerly active feature that intersects the EPR

Ž .at 348S Smith and Sandwell, 1997 . St. Helena isadjacent to an unnamed fracture zone that intersectsthe MAR at 178S. The mid-oceanic fracture zonevalleys have an added feature that tends to confirmthe proposed association with pelagic sapropels; theyare typically remote from sources of continentalsediment, a constraint that accounts for reported lowRbrNb and KrNb ratios, e.g. in HIMU basalts from

Ž .St. Helena Weaver, 1991 .

2.3.5. Examples of source mixingIt is proposed that the end-member components of

a particular oceanic island volcanic system can bejudged by the geometric relations of the trend lineson an isotope ratio plot such as Fig. 16. Examples ofdeduced single-stage mixing are provided by theWalvis and Pitcairn trends, most simply attributed tomixtures of F and MP, by the Marquesas and Tutuilatrends, attributable to mixture of F with averagesediment, and by the Rurutu, St. Helena and Tubuaitrends, apparently due to mixtures of F with HIMU.In other cases, two-stage mixing is evident wheretrends do not radiate directly from F but branch off


from another isotopic trend, presumably of priororigin. An example is provided by Terceira in theAzores. Sao Miguel is 200 km closer to the Por-tuguese and African coasts, consistent with the geo-metric deduction that its long trend line reflects amixture of F with TC. Terceira, on the other hand, isclose to the axis of the MAR, and to the Azorestriple junction, leading one to speculate that its iso-topic trend toward HIMU, apparently a branch fromthe Sao Miguel trend, may reflect recycling of blacksapropels accumulated in the axial valley or in adeep nodal valley.

ŽIn the case of the Cape Verdes Islands Gerlach et.al., 1988; Widom et al., 1997 , the Northern Cape

Verdes have an isotopic trend essentially parallel toŽ .that of above-noted Terceira Azores trend, consis-

tent with appreciable incorporation of HIMU sedi-ment. The trend of the southern Cape Verdes, on theother hand, favors incorporation of a significant MPcomponent. In this case, one cannot rule out thepossible incorporation of Saharan dust in the recy-cled mixture. The southern Cape Verdes islands arein the wind-driven path of Saharan dust, which has a

Žrelatively high strontium isotope ratio, 0.720 Rognon.et al., 1996 . Confirmation of that possibility would

require Pb isotope data for Saharan dust, difficult toassess because of worldwide industrial and automo-tive contamination by tetraethyl lead.

A distinctive feature of the Hawaiian volcanicrocks is the isotopic contrast between two parallel

Ž .volcanic chains less than 40 km apart Fig. 15 ,referred to as the ALoaB and AKeaB trends on the

Žisland of Hawaii Kurz et al., 1996; Lassiter et al.,.1996 . Both isotopic trends are consistent with addi-

tion of the MP component to old crust, but withdifferent proportions of F and DM in the old crust,the first-stage mixing product. The eastern, AKeaB

Ž .trend Kohala–Mauna Kea–Loihi Seamount has rel-Ž .atively high FeO Lassiter et al., 1996 , and gener-

ally higher Pb isotope ratios, consistent with a mix-ture of MP with MORB of a composition close to F.The ALoaB trend, on the other hand, is consistentwith a first-stage mixing product closer to DM, withan isotopic composition matching that of Cretaceouscrust, as represented by core samples from ODP Site

Ž .843 small filled circles within the MORB fieldŽ .King et al., 1993 . Similar old crust is presumed toextend beneath Hawaii. It may not be accidental that

the Loa trend, with isotopic signatures closer to thatof DM, is represented by a line of volcanoes on the

Ž .southwestern side of Hawaii Fig. 15 , closer to thecenter of Mesozoic disturbance, the proposed sourceof the DM component. Confirmation of the proposedtwo-stage mixing process, and of the second-stagedevelopment of Hawaiian lavas from recycled Pa-cific crust, is provided by geometric relations within

Žthe Nd–Pb isotopic array Staudigel et al., 1991, Fig..2c .A third extreme composition within the Hawaiian

Žisotopic field, the Koolau lavas of Oahu AKooB of.Fig. 16 , also evident in the Pitcairn and Walvis

suites, can be attributed to an FqMP mixture witha relatively large proportion of MP. A significantprimitive mantle component in the Koolau sourcehad previously been proposed by several investiga-

Žtors White, 1985; Stille et al., 1986; Valbracht et al.,.1996 , based mainly on Pb isotope ratios that fall on

the mantle geochron, together with Sr and Nd iso-tope ratios that approach those calculated for BulkSilicate Earth. The low Pb isotope ratis of the Koolaulavas can be attributed to the dominant effect of DM,unaffected by the lead-free pelagic carbonates ofMP.

2.3.6. Osmium isotope ratiosThe 187Osr186Os ratios of MORB, leached to

remove Os-rich MnO precipitates, are in the range2

1.082–1.113, slightly higher than that of abyssalŽperidotites, 1.019–1.071 Roy-Barman and Allegre,

.1994 . Plots of the osmium data in Pb–Sr isotopicŽ .space Schiano et al., 1997 showed the existence of

two separate trends of increasing radiogenic osmium,trend A associated with increasing 206 Pbr204 Pb, trendB with decreasing 206 Pbr204 Pb. The B isotopic trend

Žis a characteristic of most Hawaiian basalts Hauri etal., 1996; Bennett et al., 1996; Lassiter and Hauri,

.1998 . Both the A and B trends diverge from acommon end-member, within the field of old Pacificcrust, and defined by 87Srr86 Sr s 0.7024 and206 204 Ž .Pbr Pbs18.1 Fig. 16 . The other principalcomponent, in each case, is indicated by the respec-tive trend lines: trend A is toward a mixture of F andHIMU, trend B toward MP, attributed to mixingbetween depleted lower mantle and Mesozoic pelagicsediments. There are indications, from analyses of

Ž .the Haleakala lavas of Maui Martin et al., 1994 ,


Ž .and of lavas from Koolau Lassiter and Hauri, 1998 ,that old crust and pelagic sediment, here equatedwith MP, may be significant components in deeperzones of the Hawaiian magma source. Most of theMaui lavas have 187Osr186Os in the range 1.10"0.02, but higher ratios are observed in two relativelyyoung samples from the section marked by an abruptshift to deep-source alkalic lavas with relatively lowSr isotope ratios.

The results are consistent with prior proposalsŽ .e.g. Lassiter and Hauri, 1998 , that the variation inosmium isotope ratios of OIB reflects mixtures, inthe magma source, of old crust and marine sedimentswith relatively unradiogenic mantle osmium. Be-cause of their high osmium isotope ratios relativelysmall proportions of sediment will have a largeeffect. Analyses of age-dated pelagic clays and met-alliferous carbonates, deduced to have seawater os-

Žmium isotope ratios Pegram et al., 1992; Ravizza,.1993; Peuker-Ehrenbrink et al., 1997 show a pro-

gressive increase of 187Osr186Os during the past 80Ma, through a range from 3 to present-day ratiosnear 8. Probable osmium hosts include oceanic man-ganese nodules or Fe–Mn crusts, and the organic-richblack sapropels, identified as the HIMU component.

The basalts of HIMU islands are typically en-riched in radiogenic osmium, expressed as high187 186 187 188 ŽOsr Os and Osr Os ratios Hauri and Hart,

.1993 , as one might expect from efficient collectionof the parental rhenium by organic carbon com-pounds of the black sapropels. HIMU basalts fromSt. Helena, S. Atlantic, have low osmium content buta wide range of high osmium isotope ratios, covering

Žmost of the known range for oceanic basalt Reis-.berg et al., 1993 . Despite the indications of high

osmium isotope ratios in HIMU islands, the evidenceindicated that those islands represent isolated specialconditions not widely observed in MORB or in theprincipal island chains. It appears, on that basis, thatFe-rich pelagic sediments, including Fe–Mn nodulesand crusts, may be a more general source of highosmium isotope ratios in oceanic basalt.

Interpretation of the osmium isotope data musttake into account that the parental rhenium is ;2.6times more abundant in MORB than in OIB. Al-lowance for that constraint, together with the non-ra-diogenic character of osmium in mantle peridotites,

Ž .led Hauri and Hart 1997 to attribute the relatively

radiogenic osmium of OIB to recycling of old oceaniccrust. That suggestion is consistent with the proposedrecycling model, in which oceanic crust is a principallong-term reservoir.

2.3.7. Mesozoic submarine basaltsA feature of many volcanic islands is an apparent

association of alkalic lavas with times of restrictedlava supply, in the earliest submarine stage and atmature stages in the development of a volcano.Alkalic rocks appear at that stage and typically forma cap on the main shield. In some cases there is alate, rejuvenated stage characterized by peralkalicnepheline-bearing lavas. Prime examples are pro-

Žvided by Hawaiian volcanoes Feigenson, 1984;Clague and Dalrymple, 1988; Rhodes and Hart,

.1995 . Similar abrupt shifts from tholeiitic to alkaliclavas occur in late stages of the volcanic record for

ŽGran Canaria in the Atlantic Hoernle and.Schmincke, 1993 . Results from the Hawaii scien-

tific drilling project show abrupt appearance of alka-lic lavas in the upper 50 m of the Mauna Kea drill

Ž .core Lassiter et al., 1996 . Clague and DalrympleŽ .1988 attributed the alkalic magmas to small de-grees of partial melting of a deep garnet-bearing

Ž .source cf. Fig. 18 . A relatively deep source isconsistent with the observation, from melting experi-ments, that an alkalic melt fraction develops at higher

Žpressures than tholeiitic melts Kushiro, 1994; Ko-.giso et al., 1998 . Generation of the alkalic magmas

at greater depth implies that they will have a greaterhead of pressure, and that difference, along withtheir relatively low viscosity, accounts for their ap-parent ability to overcome the thermal constraintsthat restrict the passage of tholeiitic magma throughrelatively cold early-stage conduits, and the pressureconstraints that restrict the rise of magma to the topof a mature volcano. A characteristic feature of thealkalic lavas of oceanic islands is that most haverelatively low Sr isotope ratios, less than 0.7033Ž .Stille et al., 1986 . That signature, together withindications of a deep source, lead to the possibilitythat they include an appreciable contribution fromunderlying Mesozoic submarine lavas, expected tohave relatively low Sr isotope ratios, a consequenceof the low ratios of the Mesozoic ocean and pelagiccarbonates. Indications of other compositional differ-ences between Mesozoic and younger oceanic crust


Ž .Humler et al., 1999 provide additional support forthe proposed recycling model and serve to accountfor some of the compositional features related todepth of magma generation in oceanic island sys-tems. Reported diagnostic features of Mesozoic crustinclude relatively high CaOrAl O , which can be2 3

attributed to recycling of the abundant carbonatesediments deposited in the shallow Mesozoic ocean,and relatively high FeO, a predictable signature ofthe recycling of weathered basalt, derived from theextensive subaerial lavas of Mesozoic oceanic re-gions.

2.3.8. Subaerial MORB and cosmogenic helium-3Interpretations of the helium isotope data are con-

troversial because of differences of opinion, amonggeochemists, regarding the relative significance ofhelium isotope sources in the Earth. Helium-4 isradiogenic, a product of the uranium and thoriumsystems, and consequently is relatively abundant inancient continental rocks and derived sediments. Anexample is provided by a helium isotope transect

Žalong the Indonesian archipelago Hilton and Craig,.1989 . Their data showed an abrupt eastward de-

crease in the 3Her4 He ratio, attributed to increasedhelium-4, the result of a change from recycling of

Ž .oceanic crust Sunda Arc to recycling of continentalŽ .crust Banda Arc .

Helium-3 is from several sources, including pri-mordial helium residual from early stages of thesolar system, and cosmogenic helium, some gener-ated in minerals and rocks subjected to prolongedcosmic ray exposure, some added to sediments by

Žthe flux of extraterrestrial particles Anderson, 1993;.Farley, 1993 . The conventional interpretation of

helium isotope ratios in volcanic rocks includes theassumption that the helium-3 is mainly primordialand that high 3Her4 He ratios are the fingerprint of a

Ždeep-source mantle plume Kurz et al., 1982a,b;.Craig and Scarsi, 1998 . The concept of a deep

mantle source is open to question because of the lackof correlation of the helium isotope ratio with Sr, Nd

Ž .and Pb isotope ratios Niedermann et al., 1997 . Adeep mantle source would be expected to have acharacteristic isotopic signature. On the other handcosmogenic helium-3 is a significant component inancient surface exposures and is being developed as

Ža measure of exposure chronology Kurz, 1986a,b,

Kurz et al., 1990; Craig and Poreda, 1986; Cerling.and Craig, 1994 . In the case of helium in North

Atlantic MORB, confirmation of a dominant contri-bution from cosmogenic helium-3 emerges from thepositive covariance of the 3Her4 He ratio with FeO

Ž .content Fig. 13 , consistent with independent evi-dence of former sub-aerial exposure of large areasalong the crest of the MAR. The simplest explana-tion of the observed correlation is that both helium-3and FeO concentrations reflect an MORB source thatis dominated by multiple-stage recycling of sub-aerialbasalt and the products of prolonged weathering andcosmic-ray exposure. Similar helium anomalies havebeen found on the Shona topographic high of the

Ž .MAR at 51–538S Moreira et al., 1991 , and on theŽ .EPR at 268–288S Poreda et al., 1993 . FeO analyses

of MORB from those locations would be of interest,as a means of confirming the proposal that thehelium anomalies reflect recycling of subaerial basaltŽ .Sections 2.3.8 and 2.4.3 .

2.4. Cold spots and oceanic islands

The hotspot model of oceanic island chains, theconcept of a lithospheric plate migrating over aplume rising from a fixed deep-source hotspot, was

Žoriginally proposed for Pacific islands Wilson,.1963 . Other models have been proposed, including

Ž .the Ahot lineB suggestion of Richter 1973 that someisland chains may be developed above linear up-welling limbs of convective rolls in the upper man-tle. The hotspot model is widely accepted for theEmperor–Hawaiian chain, the classic example, andmany investigators have postulated a profusion ofhotspots to account for both oceanic and continentalvolcanic centers. Some doubts have been expressedregarding the multiplicity of hotspots, and thosedoubts were emphasized by the following quotation

Ž .from a humorous note by Holden and Vogt 1977 :

Our extrapolations show that there will be1,000,000 hotspots by the year 2000. We hopesomeone proves that hotspots do not exist, beforeit is too late.

The concept of mantle hotspots is subject to theconstraint that any hotspot within the convectingmantle will be a self-destructing feature. Conse-quently the hotspot and mantle plume theorists have


been forced to postulate hotspots at the corermantleboundary, below the level of convective flow. Deepmantle plumes will have large dimensions in accord

Žwith high viscosity see examples in Dziewonski et.al., 1993 . Large dimensions are also to be expected

in secondary plumes from the transition zone, as aconsequence of expanding viscous boundary layers.The plume model is thus faced with several con-straints: the geophysical anomalies and the localdifferences of geochemical signatures of oceanic is-lands cannot easily be accounted for by decompres-sion melting of large plumes. The following para-graphs include detailed evidence contrary to thehotspot model and consistent with an alternative

Ž .hypothesis Keith, 1993 , that oceanic island volcan-ism occurs at sites of subsidence where hydrouscrustal materials are recycled into the upper mantle.The proposed model follows one aspect of the sug-

Ž .gestion of Shaw and Jackson 1973 that early-stagemelting episodes leave a depleted residue at depth, aAgravitational anchorB that promotes continued in-flow of low-melting crustal materials. In the pro-posed model, the initial melting is not due to inter-plate friction, as suggested by Shaw and Jackson, butis a feature of crustal recycling at convection cellboundaries of two types: ocean ridge axial zones andconvective roll boundaries marked by fracture zones.Multiple episodes of partial melting and melt extrac-tion at either of these sites will leave a deep linear

Žresidue that would be relatively viscous depleted in. Žvolatiles , as well as relatively cold depleted in

.radioactive elements , and consequently would tendto remain fixed, isolated from convective flow buttending to focus continued downflow. The followingparagraphs provide a basis for making an informedchoice between two models of oceanic island volcan-ism: upwelling plumes versus focused downflow andrecycling.

2.4.1. Geophysical clues to ocean island dynamicsHawaii and the Hawaiian swell are presumed, in

current models, to represent the present location ofthe hotspot and upwelling plume that produced thechain. That concept must be questioned, however,because detailed mid-Pacific heat flow surveys showthat there is no positive thermal anomaly across the

ŽHawaiian swell Von Herzen et al., 1989; Woods et.al., 1991 . The apparent lack of a positive heat flow

anomaly is not restricted to the Hawaiian swell. SteinŽ .and Stein 1993 noted a similar deficiency of

plume-scale heat flow anomalies over the Bermuda,Cape Verde and Crozet AhotspotsB swells, and ac-cordingly favored a mainly dynamic origin rather

Žthan a thermal origin for those swells. cf. Courtney.and White, 1986 . Where heat flow data are avail-

able for exposed AhotspotB sites, Iceland and theAzores for example, high heat flow is recorded atstations near active volcanoes, but sites remote fromthe active regions typically do not show a regional

Žheat flow anomaly Palmason et al., 1979, Fig. 17;.Stein and Stein, 1993 .

Further evidence favoring the proposed recyclingmodel of oceanic island volcanism, and contrary tothe plume concept, is provided by high-resolutionseismic S-wave tomography of the central PacificŽ .Katzman et al., 1998a,b . They showed that regionalswells associated with the Hawaiian, Marquesas andSociety islands are not hot spots or hot lines; on thecontrary, the island chains are underlain by broadareas of high shear-wave velocity that extend down-ward into the transition zone and are strongly corre-lated with elevated topography and positive geoidanomalies. The results led the investigators, Are-luctantlyB, to question the hotspot hypothesis and toconsider a model in which regional swells and islandchains overlie the downwelling arms of convectiverolls. That model is opposite, of course, to the Ahot

Ž .lineB hypothesis of Richter 1973 , who proposedthat oceanic island chains overlie the upwelling armsof convective rolls.

Ž .It is suggested that the S-wave results abovecould usefully be supplemented by a comparison ofP-wave velocities for near-vertical paths through up-per mantle below a volcanic center and below sur-rounding non-volcanic regions. The proposed recy-cling model of oceanic island volcanism leads one toexpect a sub-volcanic zone of mixing and meltingbordered by downflowing mantle and underlain by adeep high-velocity zone of relatively cold residue

Ž .from repeated melting episodes cf. Fig. 4 . TheŽP-velocity test has been made at the MAR Black-

. Žman et al., 1993 and at the EPR MELT seismic.team, 1998 , and downflow is indicated by high

P-wave velocities at both sites. There is one reportedexample of an equivalent P-velocity test at a conti-nental volcanic center, the Silent Canyon caldera and


Žvolcanic center in southern Nevada W.J. Spence,.unpublished PhD thesis, Penn State Univ., 1973 .

The Silent Canyon test was unique, made possiblebecause of unusual conditions provided by a seriesof large nuclear explosions at the Nevada test site,conditions that favored analysis of sub-source seis-mic velocities and a comparison of sub-caldera man-tle with mantle up to 30 km outside of the volcaniccenter. Teleseismic data from remote seismometersindicated the presence, below the volcanic center, of

Ž .a high velocity zone HVZ which Spence suggestedmay be a mass of material that is residual from priorepisodes of melting and mantle differentiation. Henoted the trade-off between seismic velocity anddepth range, when modeling the anomalous high-velocity mass, and estimated, for example, that if theanomalous mass extends to 180 km depth, its re-quired velocity would be in a reasonable range,

Žabout 8.2 to 8.3 kmrs cf. 7.7 to 7.8 kmrs in.surrounding mantle . Some of the data favor a high-

velocity mass extending to depths approaching 400km. It seems apparent that the Nevada test-site dataare consistent with the local downflow model, orwith a seismically fast residuum, and are contrary tothe upwelling plume model of volcanism.

Several prior investigators raised questions re-Žgarding the deep plume model. Richards et al. 1988,

.p. 7691 referred to the conspicuous long wavelengthgeoid high over Hawaii, Athe classic hotspotB and

w xnoted op. cit., p. 7705 , that it is very difficult todevise a model that will produce a large positivegeoid anomaly over a plume without violating rea-sonable limits on heat flow. Several investigatorshave attempted to explain the AhotspotB –geoid para-dox by suggesting that flow-induced upwarp of inter-nal boundaries and of the Earth’s surface over arising plume, overwhelms the effect of interior den-sity contrasts driving the flow and thereby deter-mines the positive sign of the geoid anomalyŽTanimoto and Anderson, 1984; Hager et al., 1985;

.Dziewonski et al., 1993 . The upwarp model is opento question on the grounds that it neglects both theheat flow data and the classical experimental evi-dence regarding mantle traction, which showsstretching and thinning of the surface layer over a

Ž .zone of upwelling and divergent flow Griggs, 1939 .Within the broad zone of high seismic velocities

Ž .see above oceanic island systems commonly have a

narrow low-velocity zone that most investigators at-Ž .tribute to a mantle plume. Wolfe et al. 1997 used

teleseismic P- and S-wave data from Icelandic sta-tions to delineate an upper mantle low-velocity zonewith a maximal S-wave retardation of 4%, and P-wave retardation of 2%. The authors attributed theanomaly to a mantle plume but they describe thelow-velocity mass as a downward-broadening conethat is estimated to extend from 100 km to at least

Ž .400 km depth. Keller et al. 1997 pointed out that adownward-broadening cone is not the shape ex-pected of a plume head, and noted that the seismicdata can be interpreted to imply a shallow massrather than a deep plume. It seems significant that ashallow, downward-broadening cone is the geometryto be expected for the proposed recycling model, inwhich the sub-island LVZ anomaly is attributed tolocal subduction and crust:mantle mixing, with resul-tant serpentinization of mantle peridotite. The experi-

Ž .ments of Christensen 1966, Fig. 5 , show that 5% or6% of serpentine in peridotite will yield the observedS- and P-wave retardations.

The recycling model is favored over the plumemodel for the sub-Icelandic low-velocity anomaly,on the grounds that deflections of the 410 and 660seismic discontinuities are not those predicted for adeep-source mantle plume, i.e. downward deflectionof the A410B and upward deflection of the A660BŽ . ŽSection 1.1.13 cf. Bina and Helffrich, 1994; Shen

.et al., 1996 .Oceanic islands typically have a deep underlying

ŽHVZ below the zone of mixing and melting Zhang.and Tanimoto, 1993; Anderson et al., 1992 , possibly

due to basalt-depleted residues from repeated partialmelting episodes. Anderson et al. reported that vol-canic islands with HVZs in the depth range 200–300km include St. Helena, Tristan, Iceland, Easter Islandand Bouvet. Those with deeper HVZs include Azores,Ascension, Galapagos, Marquesas, Crozet and theCarolines. Hawaii was not included in the list ofthose with a deep HVZ but an earlier seismic investi-

Ž .gation Kanasewich et al. 1973 provided indicationsof a possible sub-Hawaiian HVZ at the core–mantleboundary. There is an apparent need for confirmationof that possibility.

Several other lines of evidence favor local down-flow below the Hawaiian islands, the most thor-oughly studied of all island groups. The proposed


sub-volcanic downflow is consistent with the mea-sured subsidence of Mauna Loa, recently about 2.5mmryear, slightly greater than the increased vol-canic loading, and with the seismic evidence ofextreme downward flexure, for example below OahuŽ . Ž .Fig. 17 Watts and ten Brink, 1989 . Seismic esti-mates of depth to sub-Hawaiian moho indicate aregional depression much greater than predicted by

Ž .flexure models Yingping et al., 1992; Wessel, 1993 .

2.4.2. Sub-island recycling and meltingBelow the broad HVZ typical island volcanic

centers have a central topographic depression and anunderlying LVZ that extends to about 200 km depth.Seismic data favor a model in which partial melting

Ž .occurs within the LVZ. Bock 1991 analyzed steepS–P converted seismic waves recorded in Hawaiiand concluded that the central LVZ contains a nar-row first-order seismic discontinuity at about 75 kmdepth that can be attributed to partial melting. Pres-sure and temperature conditions at that depth matchthose of hydrous melting at the dehydration depth of

Ž .hornblende Fig. 12 . The deduced sub-Hawaiianmelt zone has small areal dimensions, a feature that

Žis consistent with the proposed recycling model Fig..18 and not easily adapted to the hypothesis of

decompression melting of a large, deep-source man-tle plume.

A relevant investigation of basaltic magma gener-ation was carried out for the Azores volcanic center

Ž .in the Atlantic. Bonatti 1990 showed, by applica-tion of geochemical indicators and geothermometersto mid-Atlantic Ridge basalts and peridotites, thatupper mantle of the Azores AhotspotB has undergonea relatively high degree of partial melting but thatequilibration temperatures are the same or lower thanthose of other segments of the MAR. By comparisonwith the trace element pattern of spinel lherzolite,taken as parental to AnormalB MAR basalt, the traceelement concentrations of Azores basalts were shownto favor derivation from hornblende peridotite, simi-lar to metasomatized mantle samples from the RedSea and from St. Peter–Paul islet on the equatorialMAR. Bonatti concluded that the Azores regionappears to be a site of local volatile enrichmentrather than a AhotspotB and suggested that otherso-called hotspots may be of similar character, per-haps best characterized as AwetspotsB. That sugges-tion is consistent with the volatile-promoted meltingof the proposed recycling model.

Ž .Key evidence was provided by Hauri 1996 . Heshowed that shield-stage lavas from Hawaiian volca-

ŽFig. 17. Seismic profile through Oahu to show subsidence of the island mass under volcanic loading redrawn from Watts and ten Brink,.1989; Wessel, 1993 .


Fig. 18. Unscaled hypothetical diagram to show proposed sub-Hawaiian structure, including mantle downflow and crustal recycling above aŽ . Ž . Ž .deep residual mass R that is relatively cold depleted in radioactivity and relatively viscous depleted in volatiles . Zones of

volatile-promoted melting are focused at the boundaries between upper mantle and a subduction wedge of recycled crustal material andŽ .crust:mantle mixtures. A deep source of alkalic magma Clague and Dalrymple, 1988 is indicated schematically by a single diapir, and a

Ž .shallow level of magma accumulation is drawn at a downwarped boundary equated with a level of neutral stress cf. Fig. 6 .Ž .L and KsHawaiian Loa and Kea trends, about 40 km apart Fig. 15 , to show a proposed relationship between their distance of separation

Ž .and the width of the subduction wedge at the depth of volatile-promoted partial melting cf. Fig. 12 .

noes exhibit correlated variations in isotope ratiosand major element composition that can best beexplained by recycling of oceanic crust to the magmasource. Despite confirmation of local isotopic differ-ences, for example the distinctive Koolau, Loihi andKea trends, he favored the conventional plume modeland the existence of distinctive mantle reservoirs that

have somehow survived in the convecting mantle forperiods up to several billion years. I submit thatisolated long-term mantle reservoirs are unnecessaryas well as unlikely. The observed geochemical varia-tions of OIB can more simply be attributed to localrecycling of a single variable reservoir comprised ofupper mantle plus upper and lower oceanic crust and


intercalated sediments. The proposed recycling modelis consistent with recent melting experiments over a

Ž .range of high pressures Kogiso et al., 1998 , whichshowed that the major element diversity of OIB canbe produced from a heterogeneous crust: mantlemixture dominated by crustal components. The iso-topic signatures reflect the character of intercalatedsediments and the relative proportion of differentclasses of basalt: tholeiitic versus alkalic, sub-aerialversus sub-oceanic.

A significant feature of Hawaiian volcanism is theexistence of two separate lines of volcanoes, the

ŽALoaB and AKeaB trends, about 40 km apart Fig..15 . It is proposed, as shown schematically in Fig.

18, that they are the surface expressions of twinzones of volatile-promoted melting at the outerboundaries of a subduction wedge and that their 40km separation reflects the width of that wedge at the

effective depth of dehydration melting, ca. 75 kmŽ .Fig. 12 . The subduction wedge is proposed toinclude local oceanic crust and an underlying zone ofcrust:mantle mixing, equated with the seismic LVZ.It is tentatively suggested that similar twin zones ofdehydration melting may account for the dual vol-canic zones of southern Iceland, with axes about 100

Ž .km apart Fig. 5 . In that case the subduction wedgemay be thickened at the Icelandic tectonic bend,

Ž .from northeasterly Reykjanes Ridge to northerly innorthern Iceland.

2.4.3. Sources of high 3Her4He ratiosŽ .A recent abstract Craig and Scarsi, 1998 sum-

marized the distribution of volcanic regions withhigh 3Her4 He ratios, which they designated Ahighhelium-3 hot spotsB. The list includes the sub-aerialvolcanic regions of Yosemite and the Ethiopean rift

Fig. 19. Helium isotope ratios, 3Her4 He, of Hawaiian volcanic rocks plotted against relative age of the volcanoes. Redrawn from Kurz etŽ . Ž . Ž . Ž .al. 1983 , with blocks extended to include data from Kurz et al. 1982a,b, 1987, 1995, 1996 , Kurz 1993 , Rison and Craig 1983 , and

Ž .Kurz and Kammer 1991 . Nsnumber of analyses, RrR shelium isotope ratio of sample relative to the atmospheric ratio.A


Ž .Marty et al., 1996 , Reykjanes Ridge and Iceland,where there is independent evidence of prolongedsub-aerial exposure, and a group of islands in theequatorial Pacific. Craig and Scarsi noted a broadregional association of the Pacific Ahigh helium-3Bislands with large lower-mantle plumes delineated by

Žseismic S-wave tomography Dziewonski et al.,.1993 , proposed a genetic connection between the

plumes and the high 3Her4 He ratios, and suggestedthat the association confirms the postulated lowermantle source of primordial helium-3. The plumesare very large, however, too large to account for thedistinctive helium isotopic signatures of individual

Ž . Žislands and volcanoes Figs. 19 and 20 cf. Kurz et.al., 1983, 1987 . An additional paradox is that the

Fig. 20. Within-volcano age-related sequences of helium isotoperatios for Hawaiian volcanoes Mauna Kea and Mauna Loa. Re-

Ž .drawn from Fig. 2 of Kurz and Kammer 1991 and Fig. 4 of KurzŽ .et al. 1996 . The Mauna Kea samples are related to depth in a

research borehole, whereas the younger Mauna Loa samples areassigned approximate ages based on surface stratigraphy.

helium-3 content of most Pacific islands is one tofour orders of magnitude less than that of typical

Ž .MORB Anderson, 1998; Farley and Neroda, 1998 ,the opposite of what one would expect from thestandard model, in which MORB magmas are at-tributed to upper mantle and OIB magmas to a

Žrelatively undegassed lower mantle source Farley et.al., 1992 .

An alternative model can be based on the occur-rence of islands with high 3Her4 He in a disturbedmid-Pacific region characterized by sparse terrige-

Ž 4 .nous sediment and derived He , and by formerlywidespread voluminous sub-aerial volcanism, includ-ing numerous mid-ocean volcanoes that eventuallysubsided and produced the extensive array of atolls

Žcapped by Cretaceous reefs Hamilton, 1956;.Menard, 1964; Keating et al., 1987; Winterer, 1998 .

It is proposed that the present-day lower mantleŽ .plume Dziewonski et al., 1993 is residual from a

major Mesozoic disturbance of steady-state convec-tion, centered in the western equatorial Pacific, adisruption that produced a relatively deep boundarylayer, as indicated by the large dimensions of con-vective rolls, with wavelengths of more than 1000

Ž .km Smith and Sandwell, 1997 . Secondary roll con-vection on that scale would involve deep recyclingof oceanic crust, including the sub-aerial lavas ofMesozoic mid-Pacific, thus producing a major reser-voir of cosmogenic helium-3 in the upper mantle.

The relative importance of primordial and cosmo-genic helium-3, and of prior models favoring adeep-mantle reservoir of primordial helium-3, can beevaluated by plotting helium isotopic ratios withinisotopic space defined by lithophile elements. Thecurrent models favoring a universal high helium-3

Ž .reservoir in the mantle Patterson et al., 1994 ,Ž . Žequivalent to PHEM Primitive Helium Mantle Far-

. Ž .ley et al., 1992 , or FOZO Hart et al., 1992 , arefaced with the unsolved puzzle of opposite helium–lead isotopic correlations: a positive correlation of3Her4 He and 206 Pbr204 Pb for part of the EPRŽ .Poreda et al., 1993 , a negative correlation for At-

Žlantic MORB from 148N and 338S Hanan and Gra-.ham, 1996 . The relationship between helium and

Žstrontium isotope ratios is equally ambiguous Frey. Ž .et al., 1994 . Roden et al. 1994 concluded that

there is no significant correlation. The lack of aunique isotopic signature for high 3Her4 He islands


has led to the suggestion that the source of helium-3Žis different from that of other trace elements Graham

.et al., 1993 and that the helium isotope ratios aredecoupled from those of other isotope ratios of basaltŽ .Eiler et al., 1998 .

Ž .The opposite helium–lead correlations above ,clearly contrary to the hypothesis of a universaldeep-mantle reservoir of helium-3, can be accountedfor by the recycling model, and by the differencesbetween the Atlantic and Pacific recycling mixtures.A distinctive feature of the Atlantic crust is therelative abundance of intercalated terrigenous clasticsediments. Recycling of a crustal mixture of sub-

Ž .aerial basalt source of cosmogenic helium-3 , andsub-oceanic basalt, with intercalated terrigenous sed-

Ž .iment source of radiogenic Pb will yield a positivecorrelation of 206 Pb and 4 He, and the observed nega-tive correlation of 3Her4 He and 206 Pbr204 Pb.

The mid-Pacific recycling mix differs in that theŽ .old crust F , subject to recycling, includes a higher

proportion of subaerial basalt but a relatively minorand nearly uniform content of terrigenous sediments.As a result the recycling mixtures are dominated by

Ž .FqMP or DM Fig. 16 . The addition of either MPor DM will yield relatively low 3He and 206 Pb andproduce the observed positive correlation of 3Her4

He and 206 Pbr204 Pb. The deduced Pacific versusAtlantic contrast in terrigenous sediment componentis consistent with the relatively low 40Arr36Ar ratiosin mid-Pacific basalts, for example in Loihi samplesŽ . Ž300–2500 Hiyagon et al., 1992; Valbracht et al.,

.1997 , compared with Atlantic basalts from 208S toŽ40 36 . Ž378N Arr Ars1000–25000 Staudacher et al.,.1989 , a region in which continental sediments are

more abundantly available for recycling.

2.4.4. The helium–neon associationThe above helium data tend to favor a significant

cosmogenic source of the helium-3 in oceanic basalt,and emphasize the necessity of accounting for anapparent association between helium and neon iso-tope ratios Several investigators have utilized indi-rect correlations between helium and neon isotopicratios and interpreted the results to imply the exis-tence of a reservoir of noble gases, includingAprimordialB helium-3 and neon-20, in undegassed

Žor less degassed lower mantle Allegre et al., 1986,1993; Kurz et al., 1990; Valbracht et al., 1997; Sarda

. Ž .et al., 1988 . Hiyagon et al. 1992 and Honda et al.Ž .1993a rationalized the He–Ne systematics by as-suming that the postulated undegassed reservoir hassolar-like helium and neon isotope ratios, modifiedby time-integrated addition of a fixed proportion ofradiogenic helium-4 and nucleogenic neon-21. All ofthe above interpretations are based on the idea thatneon-20 must be primordial because of the assump-tion that there is no known nuclear reaction that

20 Žyields significant amounts of Ne Farley and.Neroda, 1998 . They noted the attendant difficulty of

explaining the observed 20 Ner22 Ne enrichments.Other studies and some of the uncertainties of inter-

Ž .pretation were brought out by Fisher 1989 . Patter-Ž .son et al. 1994 emphasized that the neon isotopic

system is the only one for which it is possible todeconvolve multiple components, because it is a

Žthree-isotope system, including neon-20 presumed. Ž .primordial and neon-21 dominantly nucleogenic ,

and the likely end-member components have charac-Ž .teristic isotope ratios. Honda et al. 1993b and

Ž .Patterson et al. 1994 noted that one can calculatean approximate 3Her4 He ratio from the 21 Ner22 Ne

21 22 Žratio by assuming that increased Ner Ne above.the solar ratio, 0.032 is due to added nucleogenic

neon-21, and that there is a concomitant addition ofradiogenic helium-4. Their mixing model is reason-able, based on the above assumptions, but does notaccount for a major paradox: the critical lack of adirect correlation between 3Her4 He and 20 Ner22 Ne.It seems likely that if the helium-3 and neon-20 areprimordial, they would have been thoroughly ho-mogenized by lower mantle convection and diffusionover the whole span of geologic time, and would beexpected to show a consistent positive correlation of3Her4 He with 20 Ner22 Ne. That is clearly not thecase; a direct plot of 3Her4 He versus 20 Ner22 NeŽ .Hiyagon et al., 1992, Fig. 7 , shows separate popu-

Ž .lations for MORB low helium isotope ratios andŽ .Loihi seamount high helium isotope ratios , but

neither group exhibits a correlation with 20 Ner22 Ne.Only 8 of the 35 analysed Loihi–Kilauea samples

Žshow neon isotopic anomalies Honda et al., 1993a,.p. 864 .

Observed correlations among isotope ratios ofneon, and with other noble gases, are not easilyadapted to the concept of a primordial source ofneon-20. 20 Ner22 Ne shows a positive correlation


21 22 Ž .with Ner Ne Honda et al., 1993c , not to beexpected if neon-20 is primordial and neon-21 is

18 Ž . Žnucleogenic, derived from O Wetherill, 1954 cf..Fig. 21 . In addition, there is a rough positive corre-

lation of the neon isotope ratios with 40Arr36Ar inŽLoihi and MORB samples Hiyagon et al., 1992;

.Valbracht et al., 1997 . The conventional model doesnot provide an obvious reason for the observedcorrelation between argon-40 and neon-20. The ob-served correlations, brought out by plotting the20 Ner22 Ne versus 40Arr36Ar data of Hiyagon et al.Ž .1992, Table 2, 8008 fraction and Valbracht et al.Ž .1997, Table 1, 10008 fraction , shows three differenttrend lines, two of which pass through the air ratio.

The above correlations favor the hypothesis thatneon-20 is not dominantly primordial but is formedby more than one process, one of which must benucleogenic. This leads to a re-examination of nu-clear reactions that might yield significant amounts

17 Žof neon-20. A source based on the reaction O a ,.20 Ž .n Ne had been proposed Kyser and Rison, 1992 ,

18 Ž .21analogous to the neon-21 reaction O a , n Ne,Ž . 17Wetherill, 1954 , but O is rare and the calculated

20 Žproduction rates of Ne in crust and mantle Kyser

Fig. 21. Systematic array of isotopes relevant to the proposednucleogenic sources of neon-20.

.and Rison, 1992 are an order of magnitude less than21 18 24 Žthat of Ne from the O reaction and from Mg n,

.21a Ne. It turns out, however, that neon-20 is alsoproduced by a higher-yielding nuclear reaction,23 Ž .20 Ž y.20Na n, a F b Ne that seems to have beenmissed in prior reviews, presumably because thefirst-stage product is 20 F. The fluorine-20 is unstableŽ . Žhalf-lifes11 s and beta decays to neon-20 Fig..21 . Another potential source of neon-20, presum-

19 Žably less productive, is the two-stage reaction F n,.20 Ž y.20

g F b Ne.A comparison of nuclear production rates of neon

isotopes was provided by my colleague, W.A. Jester,professor of nuclear engineering. The effective fast-neutron cross-section of 23 Na is slightly less thanhalf that of 24 Mg. That difference is offset, in conti-nental rocks, by the greater abundance of sodium,

Žroughly 1.5 times that of magnesium Kyser and.Rison, 1992, Table 9 , and by the difference in

isotope abundances: 23 Na constitutes 100% of natu-ral sodium, whereas 24 Mg is only 78.7% of natural

Žmagnesium. In typical granites Tuttle and Bowen,.1958, pp. 81, 114 , sodium is about 17 times as

abundant as magnesium.The recognition of a significant nucleogenic

source of neon-20 serves to account for the previ-Ž .ously puzzling noble gas correlations above , and

favors the hypothesis that the noble gas inventory ofŽoceanic basalt, except for helium-3 dominantly cos-

.mogenic? , is mainly generated in continental rocksand their weathering products, transported by conti-nent-derived sediments and added to the basalt sourceby their recycling into the melt zone. The collectionof heavy noble gases with typical atmospheric ratioscan be attributed to their trapping in interlayer sitesof clay minerals in soil profiles, while the radiogenicand nucleogenic noble gas isotopes can be attributedpartly to their generation from parent elements withinthe clays, and partly to their retention in detritalsediments derived from primary igneous minerals,including those from uranium and thorium-bearinggranitic rocks. The proposed continental source hasthe advantage, relative to a mantle source, of concen-trated U, Th and 40 K, the sources of radiogenicargon and of the neutron flux for generating nucle-ogenic noble gas isotopes. Recognition of a signifi-cant nucleogenic source of neon-20 does not, ofcourse, disprove the current hypothesis of a deep


mantle source of primordial helium-3 and neon-20. Itdoes, however, bring into question one of the foun-dations of that hypothesis; the presumption thatneon-20 must be mainly primordial because of a lackof other sources.

If one accepts the proposed continental source formost of the noble gases in oceanic basalt, particu-

Žlarly convincing for the heavy noble gases Section. 20 212.2.4 , the relatively high Ner Ne ratios in Loihi

Žbasalts seem to be anomalous see Valbracht et al.,.1997, Fig. 2 , because mid-Pacific sediments avail-

able for recycling are dominated by pelagics anddeficient in continental clastics and in 23 Na, judgedto be the principal parent of nucleogenic 20 Ne. Itseems evident that there must be a supplementarysource of neon-20, dominant in the Pacific and unre-lated to continental clastics. A long-shot possibilityemerges from the observation that neutron-inducednucleogenic reactions such as those that yield neon-20

23 19 Ž .from Na and F Fig. 21 will generally havematching cosmogenic reactions, induced by neutronsgenerated by high-energy cosmic rays. It seems pos-sible that cosmogenic neon-20 is supplied to mid-Pacific magma sources by recycling of Cretaceoussalt-flat evaporites. Salt flats and hypersaline basinsmust have been common features of Cretaceousmid-Pacific shallows, in order to have generated thebrines that produced the massive Cretaceous saltbeds. Exposed salt flat sediments presumably wouldbe a significant source of neon-20 generated by theinteraction of high energy cosmic rays with 23 Na and19 Ž .F Fig. 21 . The hypothetical salt bed source ofcosmogenic neon-20 presumably can be confirmedby measuring the isotope ratios of neon in present-daysalt flats.

2.5. Pacific disruption and oceanic island chains

The PT hypothesis includes the postulate thatvolcanism can be attributed to downflow and recy-cling at the recognized subduction zones of Indone-sia and the Pacific rim, but to upwelling and decom-pression melting at mid-ocean ridges and AhotspotsB.Some of the Pacific island chains, best exemplifiedby the Hawaiian chain, are conventionally attributedto reheating of a lithospheric plate as it migrates over

Ža deep-source hotspot Wilson, 1963; Duncan and

Clague, 1985; Bercovici et al., 1989a,b; Pringle et.al., 1993 . The hotspot and moving plate model of

island chain volcanism is faced with several prob-lems, however, including the occurrence of along-

Žchain compositional changes Hart, 1988; Graham et.al., 1988 , the typical absence of Aplume scaleB heat

flow anomalies, and the absence, for example alongthe Cook–Austral chain, of a systematic sequence of

Ž .geologic ages Okal and Batista, 1987 . For chains ofthat class, lacking a systematic age sequence, severalinvestigators have proposed an association with con-vective rolls in the mantle. The development offlow-parallel mantle rolls was proposed by KeithŽ .1972 based on convection experiments, and signifi-cant sub-ocean effects were predicted, with emphasison downward entrainment of hydrous surface materi-als at roll boundaries, with consequent serpentiniza-tion at sites of shallow recycling, and volatile-promo-

wted partial melting at sites of deeper recycling op.x Ž .cit., 1992, p. 8744 . Richter 1973 noted, from

theoretical considerations, the expected developmentof mantle rolls, emphasized the consequent alterna-tion of linear zones of compression and tension atthe surface, and suggested that mid-roll zones ofmaximal tension should be preferred sites of volcan-ism and the development of oceanic island chains.The two models thus offer conflicting views of theprocess that controls linear oceanic island volcanism:roll margin downflow and recycling in the Keithmodel, mid-roll upwelling and surface extension inthe Richter model.

Later investigators have generally favored theRichter model, commonly cited as the Ahot lineBmodel, and have demonstrated an association ofsome island chains with mantle rolls expressed as

Ž .linear geoid anomalies Baudry and Kroenke, 1991 ,Žor as fracture zones and transforms Cazenave, 1992;

.Cazenave et al., 1987; Anderson, 1994 . ExamplesŽinclude the Cook–Austral chain Turner and Jarrard,

. Ž .1982 and the Tuamotos Maia and Diament, 1991 .The Ahot lineB model can be rejected, however,based on the absence of appropriate heat flowanomalies and the indication of high sub-chain S-

Ž .wave velocities Section 2.4.1 .It is suggested that the combined geophysical and

geochemical evidence brings out the inadequacy ofthe plume model, thus leaving open the problem ofaccounting for linear oceanic island chains. Two


general postulates, essential aspects of the proposedrecycling model, can be applied to the understandingof the problem. The first is that basaltic volcanism isnot the result of decompression melting of mantleperidotite; it is due to partial melting in a heteroge-neous zone where volatile-enriched surface materialsare recycled into the upper mantle. This aspect of theproposed model is analogous to the well-establishedrelationship between subduction and volcanism inthe island arcs. The second postulate is that persis-tent volcanism at a localized site is a general resultof the development, below the melting zone, of a

Ž .mass of depleted residue Figs. 4 and 18 , that isŽ .relatively cold depleted in radioactive elements and

Ž .relatively viscous depleted in volatiles . A local coldspot of that character will tend to focus continued

Ždownflow and recycling cf. Agravitational anchor B.of Shaw and Jackson, 1973 .

With these two postulates as constraints on aworking hypothesis, it is apparent that the geometriccharacter of a volcanic zone will reflect that of thelocal recycling regime. There are indications of somelocal point sites of downflow and recycling but manyof the oceanic volcanic islands form linear chainsand therefore can best be attributed to residual coldmasses below AfossilB forms of one of the two

Ž .principal types of linear recycling zones: 1 fracturezones and transforms, the boundaries of convective

Ž .rolls in the upper mantle, and 2 axial zones atformer positions of the mid-ocean ridge. Fracturezones in the least disturbed North Atlantic sector ofthe mid-ocean ridge are sites of relatively shallowrecycling, with consequent dominance of hydrationreactions and serpentinization of mantle peridotite.There are indications, however, from the wide spac-

Ž .ing of Pacific fracture zones to more than 1000 km ,that the proposed Pacific-centered Mesozoic surge ofmantle flow was characterized by large-diameterconvective rolls, and by development of a deep,roll-boundary recycling system that would favor ac-tive volcanism and accumulation of underlying coldresidues. Those linear residues, developed below theformer axial zone or fracture zones of the ancestralMid-Pacific Ridge, are now variably displaced, andare proposed to constitute sites of end-stage declin-ing volcanism of the oceanic island chains.

The Hawaiian–Emperor chain most closely fitsthe hotspot model and is the usually cited prime

example. Volcanic rocks of the Emperor seamountsand the Hawaiian chain exhibit a regular sequence of

Ž .radiometric ages, from oldest ;80 Ma at thenorthwestern end to youngest at Loihi seamount, off

Ž .Hawaii Dalrymple et al., 1980 . It is proposed, as aworking hypothesis, that the Hawaiian–Emperorchain, in contrast to most other chains, is locatedalong the trend of a linear cold residue that wasdeveloped below the ancestral Mid-Pacific RidgeŽ .Fig. 22 , and that part of that residue was leftbehind when the ridge was variably displaced duringa Mesozoic disruption of Pacific mantle convection.

The sequence of volcanic ages, uniquely system-atic, is attributed to interaction between the down-flow focusing effect of that residue and a downflowtendency at the outer, radial-flow limits of a Polyne-sian AsuperplumeB and its retreating, post-Cretaceousexpression. The Afocused downflowB model of se-quential Hawaiian–Emperor volcanism is repre-sented in Fig. 22, in a simplified form, including theextent of a large present-day plume in the lowermantle below the equatorial Pacific. That plume, as

Žoutlined by seismic S-wave tomography Dziewon-.ski et al., 1993 , extends over more than 7000 km,

from the Mariana Arc to Polynesia, where it peaksbelow the South Pacific isotopic and thermal anomalyŽ . Ž .SOPITA Staudigel et al., 1991 , approximately

Ž .centered under Tahiti T in Fig. 22 . It is proposed tobe the residual expression of a major Mesozoicplume, presumably triggered by a cluster of largemeteorite impacts. The original plume presumablyhad a whole-mantle configuration, involving outwardradial flow in the upper mantle. Within the contextof the proposed impact model, the early stages ofimpact and Pacific disruption are proposed to includerapid formation of the 25-km-thick crust at Ontong–

Ž .Java Plateau Tarduno et al., 1991 , the concentricmagnetic anomaly pattern around the East Mariana

Ž .and Pigafetta basins Castillo et al., 1994 , and thenumerous Cretaceous volcanoes that produced the

Ž .Darwin Rise seamounts Janney and Castillo, 1999 .The time of disturbance is constrained by Jurassicand Cretaceous radiometric ages of the oldest vol-

Žcanic rocks up to 167 Ma for Pigafetta Basin tholei-. Ž .ites Castillo et al., 1992 , and by paleontological

dates of the oldest sediment: 130–150 Ma for ShatskyRise, 100–125 Ma for the Ontong–Java PlateauŽ .Larson, 1991 .


Ž .Fig. 22. Pacific Island chains redrawn from Larson, 1991 , and a succession of dated circles to show the proposed contraction, 60 Ma toŽ .Recent, of a hypothetical zone of outward flow and downwelling, a radial pattern above an elongated lower mantle plume dashed outline

Ž .based on S-wave tomography Dziewonski et al., 1993 , and deduced to represent a Mesozoic impact-triggered disturbance of theŽ . Ž .steady-state convective system see text . Large Cretaceous plateaus of the impact region, with ages of oldest sediment Larson, 1991

Ž . Ž . Ž .include: Manihiki M, 115–125 Ma , Marcus–Wake Seamounts MW, 90–115 Ma , and Ontong–Java OJ, 100–125 Ma .ŽThe peak of the persistent residual plume corresponds approximately with the broad Polynesian superswell and SOPITA anomaly Staudigel

. Ž . Ž .et al., 1991 , approximately centered at Tahiti T . Final stage downflow is indicated, from interpretation of seismic anisotropy see text , toŽ .be focused on an elongated mid-Pacific zone trend-line HM , that extends from Hawaii to the Marquesas and is tentatively identified as the

residual trace of a cold, depleted mass developed below the ancestral Mid-Pacific Ridge.

The early-stage flow pattern, triggered by theproposed western Pacific impact, is proposed to havegenerated the circum-Pacific subduction zones at thepeak of the Mesozoic surge and to have formed theretreating circle of downflow in post-Cretaceous timeŽ .Fig. 22 .

In the proposed Pacific model, a portion of theaccumulated cold residue below the ancestral Mid-Pacific Ridge is presumed to have been displacedeastward, possibly represented by the deep high-velocity masses identified in the lower mantle below

Ž .the Americas Van der Hilst et al., 1997 . Theremaining linear residue presumably would have amuch weakened focusing effect. Similarly, thedownflow tendency at the retreating outer limits ofthe Pacific superplume would be a weakened resid-ual effect, appreciably less than at the Cretaceous

maximum. The essence of the proposed Hawaiian–Emperor model is that although downflow focusingprobably was relatively weak at both controllingsites during Cretaceous to Recent time, it would besignificantly enhanced at the intersection of theresidual sub-ridge cold residue with downflow at thecontracting outer limits of radial flow around thePolynesian superplume. The trace of the migratingintersection is proposed to have generated the 80 Mato present-day sequence of volcanism along theHawaiian–Emperor chain.

Evidence of a persistent trace of the cold residueis provided by a detailed study of Pacific seismic

Ž .anisotropy Ekstrom and Dziewonski, 1998 . In addi-tion to the pattern of anisotropy related to the EPRŽ .see Section 1.1.11 , the broad Pacific region remotefrom the EPR is characterized by a pattern of upper


mantle anisotropy with a mid-Pacific focus. EkstromŽ .and Dziewonski 1998 plotted the difference be-

Žtween V and V , in which V e.g. of a RayleighSV SH SV.wave is the velocity of a vertically travelling S-wave,

Ž .and V e.g. of a Love wave is the velocity of aSH

transversely polarized S-wave following a horizontalpath. They showed, for intermediate period Love-waves and Rayleigh-waves, that there is a radialpattern of V fast directions, unseen in other oceanicSH

regions, that converges toward an elongated focalzone extending from the Hawaiian chain to the Mar-

Ž .quesas dashed trend line HM in Fig. 22 , a linearzone in which the negative V –V anomaly peaksSV SH

at more than 4% at 125 km depth, appreciablydeeper than the ;60 km peak depth of the S-wave

Žanomaly below the axis of the EPR cf. Forsyth,.1992, Fig. 28 . Ekstrom and Dziewonski concluded

that the origin of the unique Pacific anisotropicanomaly is unknown, but the proposed disruptionmodel of Pacific structure, together with the centralPacific position and trend of the elongated H–Mfocal zone, leads to the possibility that the finalstages of the disruption-generated flow pattern arefocused on the linear cold zone developed below theancestral mid-Pacific Ridge. The radial pattern ofLove-wave anisotropy does not, by itself, differenti-ate between divergent and convergent flow aroundthe mid-Pacific focal zone. It is evident, however,from the identification of the Hawaiian and Marque-sas chains with a deep zone of high seismic veloci-

Ž .ties Katzman et al., 1998a,b , that they are sites ofmantle downflow, rather than plume sites. In accordwith that conclusion, the radial pattern of fast Love-waves can be attributed to convergent flow towardthe Hawaiian–Marquesas line. It is assumed that thestress regime at the effective depth of the seismic

Ž .anomaly is uniaxial compression see hex-b modelŽ .Section 1, Fig. 3 and that the fast Love-waves arepolarized in the a–c plane, and propagating parallelto the olivine b axis, and to the direction of shorten-ing.

No attempt was made to fit other Pacific islandchains into the above migrating focus model. Mostof them lie too close to the Polynesian center ofdisturbed mantle flow and would not be expected toshow a long-term sequence of volcanic ages. The

Ž .Easter-Sala y Gomez chain E–W line at 258S ,does, however, exhibit a westerly decreasing age

sequence that is roughly consistent with the proposedŽ .retreating circle of radial mantle flow Fig. 22 .

Ž .Radiometric ages O’Connor et al., 1995 range from26 Ma at the eastern end to near zero at EasterIsland.

The above summarized geophysical evidence ofoceanic island subsidence within a surrounding zoneof normal to low heat flow and high seismic veloci-ties indicates sub-island mantle downflow. The de-duced downflow regimes are confirmed by the iso-topic diversity of OIB and by geochemical signaturesthat are correlated with those of local crust andentrained sediment. The spectrum of evidence favorsgeneration of oceanic basalt by local crustal recy-cling and volatile-promoted melting and implies aneed for critical re-examination of the popular con-cept of basalt generation by decompression meltingof deep-source mantle plumes. The North Atlanticstructure appears to be the prime example of aleast-disturbed steady-state flow regime, and the Pa-cific, Caribbean and Scotia arc structures can beregarded as extreme results of disrupted flow regimesŽ .Haxby, 1985; Smith and Sandwell, 1997 .

The real voyage of discovery consists not inseeking new landscapes but in having new eyes . . .Ž .M. Proust

3. Continental rifts and ranges

3.1. Continental mountain ranges

The PT hypothesis has been dominant in geody-namics for more than two decades and was ex-panded, at an early stage, to include the plate colli-

Žsion model of mountain range formation Dewey and.Bird, 1970; Dewey et al., 1973; LeFort, 1975 . Some

reservations were expressed, along the way, regard-ing attempts to fit complex continental ranges intothe relatively simplistic plate collision model; for

Ž .example see Trumpy 1975 , but the great majorityof investigators have accepted the PT model orproposed minor modifications regarding the questionas to whether moving plates are the drive mechanism

Ž .or are driven by mantle flow Turcotte, 1982 . Aproposed viscous flow model of sub-range mantledynamics is based on seismic anisotropy, gravity


anomalies and crustal structure, in conjunction withexperimental convection and hot creep. The pro-

Ž . Ž .posed viscous flow VF model Keith, 1993 em-phasizes the weakness of large rock masses on ageological time frame and favors the concept of theuppermost mantle as a thick viscous boundary layerrather than as rigid plates. The VF model is not new;it is essentially similar to the classic model of GriggsŽ .1939 and to several proposed modifications of

Ž .plate tectonics. Houseman et al. 1981 and Platt andŽ .England 1994 followed Parsons and McKenzie

Ž .1978 in postulating that the lower part of thelithosphere below a mountain range acts as a Ather-mal boundary layerB, normally a conductive layerbut subject to secondary convection following anepisode of compressional thickening. The thermalboundary concept was adapted by Houseman et al.Ž .1981 to imply that the effective creep viscosity ofthe lithosphere decreases with depth and that near-basal lithosphere approaches the viscosity of under-

Žlying asthenosphere cf. Molnar, 1988; Willet and.Beaumont, 1994, Fig. 2d . Indications of hydrated

Ž .upper mantle Haggerty, 1997 lead one to implyŽreduced creep strength of peridotite Mei and Kohlst-

.edt, 1997 and to favor an upper mantle model basedon hot creep rather than on rigid lithospheric plates.That conclusion is consistent with the most recentstudy of strain rates and gravitational potential across

Žthe Himalaya and the Tibetan Plateau England and.Molnar, 1997 . They showed that the dynamics of

active deformation obey the equations of creepingflow and that the continental lithosphere should beregarded as part of the fluid portion of the solidEarth.

3.1.1. SKS data and interpretationsSeismic anisotropy, a principal indicator of the

direction of mantle flow, is attributed mainly topreferred orientation of olivine, the most abundantand most effectively oriented mineral of mantle peri-

Ždotites Christensen, 1966, 1984; Nicolas and Chris-.tensen, 1987; Mainprice and Silver, 1993 . In a

shear-dominated flow regime, presumably typical oflateral upper mantle flow beneath a lid, there is astatistical tendency for olivine to recrystallize orrotate into an orientation of easy deformation, withcrystallographic a parallel to the flow directionŽ . Ž .Zhang and Karato, 1995 . The preferred 010 slip

plane and slip direction, parallel to crystallographica, tend to become aligned with the foliation plane

Ž .and lineation of the rock fabric Fig. 23, Inset A .ŽThe crystallographic a axis of olivine pole to

.100 face is the direction of fastest P-wave propaga-tion. It is also the vibration direction of the fast

Ž .S-wave of a split S-wave pair Table 2 and, within aregime of shear flow, can be assumed to represent

Žthe direction of mantle flow Christensen and.Lundquist, 1982; Karato et al., 1989; Kern, 1993 .

That assumption is not valid, however, in a region ofuniaxial compression. Under conditions of uniaxialcompression, to be expected in a region of conver-gent mantle flow, interpretations are complicatedbecause the preferred orientation of olivine is 908from that developed in a shear regime. Olivine tendsto rotate or recrystallize to an orientation of difficult

Ž .deformation Ahex-bB model of Fig. 23 , in whichthe b axis is aligned with the direction of shorteningand the a and c axes form a girdle in the flattening

Žplane normal to b Ave Lallemant and Carter; Nico-´.las et al., 1973; Karato, 1987 .

Several recent investigators of mountain rangedynamics have deduced mantle flow directions fromseismic anisotropy, based on the phenomenon ofshear wave splitting in anisotropic mantle, analogousto optical splitting in anisotropic crystals. Pairs ofshear waves are formed by the splitting of teleseis-mic SKS waves generated by P-to-S conversions atthe core–mantle boundary. If the lower mantle isisotropic, the initial SKS wave becomes polarized

Žwith one wave in the AradialB orientation in the.vertical plane containing the ray path . SKS waves

typically arrive on near-vertical paths. In regions ofanisotropic mantle, an SKS wave splits into a fastS-wave polarized parallel to the preferred orientationof the olivine a axis and a slower ray polarized in a

Žplane perpendicular to a Christensen, 1966, 1984;Babuska and Cara, 1991; Mainprice and Silver,

.1993 . The direction of polarization of the fast wave,and the time delays between fast and slow yieldinformation regarding lattice-preferred orientation ofolivine. For the relatively simple case of horizontalshear flow, the vibration direction of the fast S-wavecan be equated with the preferred orientation of thea axis and with the flow direction of the uppermantle below the receiving station. Interpretationscan be based either on the orthorhombic model of


Fig. 23. Simplified hypothetical section normal to a mountain range and borderlands, to show subjacent mantle treated as a viscousboundary layer in which shear flow is dominant in a sub-lid zone, ASB, and there is a downward gradation to massive body flow. Thediagram is designed to emphasize proposed relationships of mantle flow to olivine orientation and related seismic anisotropy under

Ž . Ž .conditions of shear flow Ahex-aB model appropriate , and in the regime of uniaxial compression Ahex-bB model, see text . Inset A showsthe relationship of the simplified quasi-hexagonal, Ahex-aB model to the correct orthorhombic model of olivine, in which seismic P-wave

Ž . Ž . Ž .velocities are in the order a fast , b slow and c intermediate . The same order of velocities applies to S-waves vibrating parallel to a, b,and c. The contrasted olivine orientations are based on prior laboratory studies of anisotropy in peridotite samples and on investigations of

Žseismic velocities in ophiolites Christensen and Lundquist, 1982; Karato, 1987, Karato et al., 1989; Mueller, 1982; Borman et al., 1993;. Ž . Ž .Babuska et al., 1993 . Seismic station, diagrammatic example k . Mantle flow heavy-line arrows .

Ž . Ž .SKS and dashed line: example of a steeply propagating shear wave split into fast- filled circles and slow-waves transverse double arrowsŽ .representing vibration sub-parallel to b heavy line axis within the zone of uniaxial compression and in an underlying zone of downflow.

olivine, or on a quasi-hexagonal Ahex-aB model thatŽ .is used to simplify the interpretations Fig. 23 .

A surprising recent development in mountainrange dynamics is the publication of a class ofpapers in which it was concluded, from SKS data,that the azimuth of polarization of the fast shearwave, and the deduced direction of sub-range uppermantle flow, are horizontal and parallel to the axis ofthe associated mountain range. Various explanationshave been offered, typically that the range-parallelmantle flow is a result of Aplate collisionB and

ŽAsideways expulsionB Silver and Chan, 1991; Nico-.las, 1993 . That interpretation has now been applied

Žto many of the major mountain ranges Vauchez and.Nicolas, 1991; McNamara et al., 1994 , including the

Ž .Alps Borman et al., 1993, p. 165 , and the AndesŽ .Silver and Chan, 1991; Russo and Silver, 1994 . Ji

Ž .et al. 1994, p. 21 cite several examples of deducedŽrange-parallel mantle flow Tien Shan, Balkans,

.Caucasus, Pamir and Hindu-Kush based on range-parallel polarization of the SKS fast wave. Someinvestigators have expressed reservations or pro-

Ž .posed alternative models e.g. Borman et al., 1993 .Ž .Vinnik et al. 1992 noted that in regions of conver-

gent flow the polarization direction of the fast wavetends to align normal to the direction of shortening


Ž .cf. hex-b model of Fig. 23 . Other investigatorshave reported range-parallel polarization of the fastSKS shear wave without attributing that result to

Žrange-parallel mantle flow Savage et al., 1989;.Souriau and Njike-Kassala, 1993 . Earthquakes from

an active source at about 160 km depth below theŽ .Andes Shih et al., 1991 , showed preferred orienta-

tion of the fast S-wave parallel to the range ratherthan to the hypothetical plate motion.

A recent extreme example of deduced range-parallel flow, based on orientation of the fast S-wave,

Žis the conclusion Russo and Silver, 1994; Polet et.al., 1996 that the upper mantle flows laterally north-

ward and southward beneath the South Americantrench and the Andes from a AstagnationB juncture

Žnear the major bend of coast and mountains ca..208S , and eventually flows half way around South

America and from the Pacific to the Atlantic viaflow channels beneath the Caribbean and ScotiaSeas. The investigators initially expressed surprise at

Ž .the lateral flow results Silver and Chan, 1991 , butlater concluded that the proposed trench-parallel flowpattern could, after all, be adapted to the plate colli-

Ž .sion model Russo and Silver, 1994 . They appar-ently favored the lateral flow model in part becauseof a perceived space problem related to the platetectonics model, a need to allow for the volume ofsubducted plates and to counter a Ashrinking PacificBby a transfer of material from Pacific to Atlantic.

Recent reports of SKS-derived seismic anisotropyand deduced range-parallel mantle flow are based onthe use of a quasi-hexagonal model of velocities in

Žolivine Keith and Crampin, 1977; Montagner and.Nataf, 1988 with crystallographic a as the symme-

Ž .try axis polarization direction of the fast S-wave .That model, here designated the Ahex-aB model, isan acceptable simplification for the case of lateralshear flow of uppermost mantle beneath a lid butmay yield ambiguous results for two cases: the caseof steep upward or downward shear flow where a isnearly vertical, and the case of uniaxial compressionnear a flow corner, where b will be the axis of

Ž .symmetry Fig. 23, hex-b, Inset . An alternativemodel of mountain range dynamics is summarizedbelow and compared with seismic data from theAlpine region. The essence of the proposed hypothe-sis is that the uppermost mantle can be treated as a

Ž .thick viscous boundary layer Keith, 1993 , in which

shear flow is dominant in a relatively thin sub-crustallayer, whereas the underlying mantle is dominatedby massive body flow. Compressional mountainranges are attributed to traction of subjacent viscousmantle flow at zones of convergence and downflow.

Ž .The proposed boundary layer features Fig. 23are similar to those of the proposed sub-oceanic

Ž .model Section 1 and the postulated sub-rangedownflow are comparable to that proposed by Laub-

Ž . Ž .scher 1974 and Mueller 1982, 1989 as a late stagein the Alpine dynamic system. The width of thecorner zone and downflow column presumably de-pends on several parameters, including viscosity and

Ž .heating mode cf. McKenzie, 1979, Fig. 7 , andcannot be defined exactly from available data. Thehypothesis of deep sub-range crustal recycling, basedmainly on seismic evidence, is supported by theevidence of resurgent serpentinized peridotite forwhich petrographic thermobarometers indicate equi-

Žlibration at depths of 70 to 150 km Ernst, 1978;.Green, 1995 . In defense of the omission of the

conventional designation of a separate lithosphere,effectively isolated from mantle convection, I notethat the concept of an isolated lithosphere is well

Žestablished for ancient continental shields Jordan,.1975; Boyd and Gurney, 1986 but cannot as clearly

be defined for active tectonic regions.

3.1.2. Sub-alpine anisotropy and deduced mantleflow

Ž .The hypothetical section Fig. 23 provides a ba-sis for a comparison of the model with seismic datafor the well-studied region adjoining the northern

Ž .boundary of the eastern Alps Fig. 24 . Seismic dataindicate that the fast P direction for central Europen

bordering the Alps, commonly equated with a axisalignment and mantle flow direction, is about N208EŽ .Bamford, 1977 , nearly orthogonal to the Alpinefront and consistent with either Aplate collisionB orthe proposed viscous flow model. In notable con-trast, the data based on splitting of steep-path SKS

Ž .waves Vinnik et al., 1994 shows the polarizationdirection of the fast S-wave nearly parallel to theadjacent Alpine front. It is proposed that the 908contrast is due to a difference in the depth rangessampled by P waves and by steep-path SKS shearn

waves. The P waves traverse the uppermost shear-n

dominated zone of the mantle, whereas the anisotro-


Fig. 24. Plan view of the region to the Northwest of the AlpineFront, to show preferred average orientation of the olivine a axis

Ž . Žat shallow depths arrow , as deduced from P-wave data Bam-.ford, 1977 , compared with preferred orientation at greater depth,

as indicated by alignment of the fast S-wave of SKS shear wavesŽ .Borman et al., 1993; Polet et al., 1996 .

py of steep-path SKS waves includes a major contri-bution from a deeper zone, dominated by uniaxialcompression, where the vibration direction of the fastwave is in the a–c plane, orthogonal to the direction

Ž .of shortening hex-b model .The validity of the Ahex-bB model and its rele-

vance to regimes of uniaxial compression at zones ofconvergent flow in the upper mantle are supportedby petrofabric studies of experimentally and natu-

Žrally deformed mantle rocks Ave Lallemant and´.Carter, 1970; Nicolas et al., 1973 , some of which

exhibit b axis concentrated normal to the foliation,and have a and c axes randomly oriented in a girdle

Žwithin the foliation plane Christensen, 1984; Ribe. Žand Yu, 1991; Babuska et al., 1993, p. 176 cf. Fig.

.23 . A similar orientation was reported by KernŽ .1993 in an upper mantle lherzolite xenolith fromthe Amassif centralB of France. The random orienta-tion of a and c in the foliation plane of that speci-men was confirmed by the absence of splitting inshear waves propagating parallel to b. It is apparentthat interpretations of upper mantle flow in mountain

regions depend on a reasoned choice between theAhex-aB and Ahex-bB models of olivine.

Below a mountain borderland, and with idealizedolivine orientations as represented in Fig. 23, anupward-propagating split SKS wave will be con-strained by a 908 difference in fast vibration direc-tions: range-parallel fast in a lower layer of uniaxial

Ž .compression hex-b model and range-normal fast inŽ .the upper, shear flow layer hex-a model . There are

indications, judging from the shallow paths of P -n

waves, that the upper layer is relatively thin, proba-bly not more than 50 km thick in continental regions.In the mantle directly below the range axis, the b

Ž .axis slow will have a range-normal orientation inthe zone of uniaxial compression and in the underly-ing zone of downward shear flow. A resultant ideal-ized split-wave path is represented by the dashed lineŽ .SKS terminating at hypothetical receiving station kof Fig. 23. It seems likely that the range-normal

Žorientation of b in those two stress regimes rela-.tively fast range-parallel vibration has led some

investigators to imply range-parallel mantle flow aspart of the plate collision model of mountain rangeformation.

Supporting evidence for the hypothesis of sub-Al-pine downflow, and contrary to the concept ofrange-parallel lateral flow, is provided by availabledata for the velocities of steep-path teleseismic P-waves, which have the advantage of relatively simple

Žgeometry and larger anisotropies than S-waves Ta-.ble 2 . In a depth zone of downflow the stress regime

will be dominated by vertical shear flow, crystallo-graphic a will be nearly vertical and steep-pathP-waves will have maximal velocity. An example is

Žprovided by the reports of Babuska et al. 1987,.1993 that steep sub-Alpine paths have the maximum

P-wave velocity in the uppermost 100 to 180 km ofmantle, faster than those below Alpine borderlands.

Ž .Mueller 1989 concluded that the P-wave data forsub-Alpine mantle indicate a cold, dense, slowlysubsiding lithospheric root that has penetrated to adepth of 130 to 200 km. Sub-Alpine downflowseems to be further indicated by the occurrence ofmasses of serpentinized peridotite in the western

Ž .Alps Ernst, 1978 . One of them, the Alpe Aramimassif, judged to have been tectonically emplaced asa diapir during terminal stages of Alpine orogeny, isa serpentinized garnet lherzolite bordered by altered


eclogite and chloritized peridotite mylonite. Petro-graphic thermobarometers from that massif yieldedequilibrium pressures equivalent to depths exceeding70 km and Aprobably approaching 150 kmB.

3.1.3. Application to other mountain rangesAdditional examples of evident sub-range down-

flow are provided by P- and S-wave data for othermountain ranges. Seismic refraction and reflectiondata indicate thicker than average crust in the south-ern Sierra Nevada Range, and high velocity sub-range

Žstructures that extend to 200–250 km depths Rup-.pert et al., 1998 . Similar anomalies are reported

below the San Bernardino Mountains, at the easternŽend of the Transverse Range Humphreys et al.,

.1984, 1989 . The authors attributed the sub-rangedense mass to downflow that is driving local tectonicactivity, including compression of the TransverseRanges, a concept that is essentially similar to the

Ž .proposed general model Fig. 23 . A summary ofseismic shear-wave splitting of SKS waves beneath

Ž .the Pyrenees Barruol et al., 1998 indicated a simi-lar stress regime; the polarization direction of thefast S-wave is parallel to the trend of the Pyreneanbelt, consistent with N–S compression and with theolivine b axis oriented in the direction of shortening.The axis of the Pyrenees is also marked by a nega-tive Bouguer gravity anomaly, attributed to abnor-mally thickened crust and by a slow seismic-velocityanomaly at 50–100 km depth, attributed to sub-range

wsubduction of thickened lower crust op. cit., 1998,xp. 30,045 . All of these features are consistent with

the Alpine data and favor the proposed model ofconvergent viscous flow and downwelling below amountain range.

Interpretations for many of the largest ranges,exemplified by the Andes, are complicated becauseof asymmetry related to their location at a conti-nent–ocean boundary. They are conventionally at-tributed to downgoing lithospheric plates, but can,within the framework of the proposed VF model, beconsidered as convection cell boundaries that havebeen displaced from a mid-oceanic to a continentalmargin position and therefore have developed anasymmetry at the confluence of relatively warmsub-continental mantle and colder sub-oceanic man-

Žtle. With that model in mind cf. England and Mol-.nar, 1997 the subduction angle can be attributed to

the thermal contrast at the confluence. The proposedrelatively warm mantle on the continental side of theconfluence is evidenced by the high heat flow andlow S-wave velocities below back-arc basins.

V data indicate a narrow root zone sub-parallelP

to the Himalayan chain, characterized by high P-velocity and extending to at least 400 km below the

Žbordering Southern Plateau Zhou et al., 1995; cf.S-wave data of Sandvol et al., 1994; gravity profiles

.of Malinconico, 1989 . The cited examples of P-velocity evidence favoring sub-range downflow donot support a general rejection of the hypothesis ofrange-parallel or trench-parallel mantle flow for othermountain systems but may encourage investigatorsto make the seismic velocity tests that are needed tochoose between it and the downwelling model.

3.2. Continental rifts

3.2.1. The rift connection: ocean to continentThe evidence for an alternative dynamic model of

mid-ocean ridges—convergence, downflow and re-Ž .cycling below the oceanic axial zone Section 1 —

leads to a re-examination of continental analoguesand to a general hypothesis: that persistent magmaticactivity occurs at sites of downwelling and recycling,not at Ahot spotsB. The axial zone can be traced, viaAfar, to the East African rift system and thence tomobile belts in southern Africa. Afar is a triplejunction, the confluence of three convection cells inthe proposed model, and is expected to be a site ofstrongly focused downflow. It should be marked by alarge volume of accumulated subduction mixture,characterized by the low seismic velocities of serpen-tinized peridotite. The connecting rifts and mobilebelts are postulated to be surface expressions ofconvection cell boundaries. They share a similartectonic setting, a characteristic association with theborders of ancient cratons, and both yield a similarvariety of igneous products, from basalts to granitesand the characteristic assemblage of carbonatites andperalkaline silicate rocks. The ocean to continentconnectivity, together with the co-linearity and tec-tonic association of rifts and mobile zones, leads oneto question the current extensional hypothesis ofrifting and support the proposition that oceanic andcontinental rifts are surface expressions of a commondrive system: convergence and downflow at upper


mantle convection cell boundaries. This paper em-phasizes the structural and geochemical evidence fordownflow and recycling below continental rifts, us-ing the African rifts as prime examples.

Recent investigations of sub-African mantle dy-namics favor upwelling beneath the ancient Archeancratons, a feature that is in accord with experimentalevidence of preferred upwelling beneath a thermal

Ž .barrier Section 1.1.2 . An example is provided bythe heat flow pattern in the South African Kaapvaal

Ž .craton and surroundings Ballard and Pollack, 1987 ,interpreted to imply sub-craton upwelling and deflec-tion of rising mantle flow from sub-craton keel to thebase of surrounding mobile zones. The existence of athick sub-craton keel is confirmed by petrologic datathat yield an inflected geotherm in kimberlite-rafted

Ž .mantle xenoliths at 175 to 250 km , a depth thatcorresponds to a change from equigranular textureŽ . Žsub-craton keel to sheared flow textures astheno-

. Ž .sphere Boyd and Gurney, 1986; Keith, 1993 . AŽsimilar heat flow pattern, from low values avg. 40

2 . 2mWrm in mid-craton to 70 mWrm beyond cra-ton borders, was subsequently reported for the Tan-

Žzania craton of equatorial Africa Nyblade et al.,.1990; Nyblade and Pollack, 1993 , and deduced to

be a typical feature of ancient cratons and theirsurroundings. The existence of a broad thermal up-welling beneath Tanzania was recently confirmed bydata from global network seismic stations supple-mented by 20 broadband stations in TanzaniaŽ .Ritsema et al., 1998 . The Tanzania seismic experi-ments also show that sub-rift earthquakes have depths

Ž .as great as 45 km Rukwa graben , attributed toŽ .faulting in uppermost mantle Langston et al., 1998 .

Brittle behavior at that depth seems contrary to theconcept of sub-rift mantle upwelling and decompres-sion melting.

3.2.2. Proposed rift modelThe indications of relatively dense sinking mantle

Žbelow the axial valley of the mid-ocean rift Sections.1.1.10–1.1.12 , supports the proposal that continen-

tal rifts, like their oceanic analogues, are collapsestructures, surface expressions of convergence and

Ždownflow cf. original rift definition, Gary et al.,.1965 . Some of the African rifts can be traced by

fault patterns and seismicity to older mobile belts,and some rifts are clearly within those belts and

related to reactivation of tectonic mobility at cratonborders. Examples are provided by the Early Creta-ceous rifts of Western and Central Africa, shown tohave developed within the system of Pan-African

Žmobile belts that were active around 600 Ma Maurin.and Guiraud, 1993 . In the case of typical mobile

zones, a relationship to convergent flow and com-pression is not seriously in dispute; they clearlyinvolve tectonic shortening. The mobile belts, roots

Ž .of former orogenic belts Clifford, 1966 are zonesof repeated tectonic activity, commonly compres-

Ž .sional Colliston et al., 1991 , characterized by tightfolds and thrust structures in high grade gneiss ter-

Ž .ranes Ring, 1993 . Examples are provided by thestructure of mobile zones bordering the Kaapvaalcraton and by the Damara mobile zone, between the

ŽCongo and Kalahari cratons of SW Africa Tankard.et al., 1982, pp. 221–228, 315, 318 . There are many

complications, including granitization of the de-formed gneisses, and the development of extensivesub-vertical shear zones, some as long as 600 kmparallel to the belt. To the south of the clearlydefined Eastern and Western rift systems, they cometogether in the Malawi rift and thence splay out toform a wide tectonic system of rifting and volcanicactivity, mainly along the eastern border of the Tan-

Ž .zania craton Ebinger et al., 1987, 1997 . Volcaniccenters follow the transition zone between Archeanand Proterozoic regions or lie within the borderingProterozoic mobile zone. The common tectonic set-ting of rifts and mobile zones, at craton marginsŽ .Fig. 25 , leads to the question as to whether themobile zones are deep expressions and the riftsshallower expressions of regions of convergence anddownflow in underlying mantle. The strong compres-sional features of mobile zones favor a cell-boundaryconvergent flow regime for mobile zones and relatedrifts. A seismic reflection profile across the Malawi

Ž .rift at 108S Ebinger et al., 1987 shows downward-increasing dips of valley-fill sediments and a com-pressional fold regime at depth.

3.2.3. Structural and geophysical eÕidenceEpisodic reactivation of many of the rifts and

mobile belts, in widely separate time frames, isevidence of a persistent but variable controllingmechanism. An obvious candidate is focused mantleflow, accelerated during orogenic episodes, associ-


Fig. 25. Distribution of African carbonatite complexes in relation to the margins of Archean cratons and to bordering rift valleys and mobileŽ .zones carbonatites after Wooley, 1989; cratons after Kampunzu and Popoff, 1991 . The indicated cratons, characterized by thick

Ž . Ž . Ž .lithosphere, and by tectonic stability for more than 2000 Ma, include the following N to S : West African WA , Uweinat Remnant U ,Ž . Ž . Ž .Congo C , Tanzania T and Kalahari K , the last commonly mapped as two separate cratons, the Zimbabwe and Kaapvaal Cratons,

northern and southern segments divided by the NE-trending Limpopo Mobile Belt. Carbonatites are notably associated with the East AfricanŽ . Ž .rift system, with the Atlantic coast of southern Africa see text and with the following: Mozambique belt eastern boundary of T and K

Ž . ŽRing, 1993; Nyblade and Pollack, 1993 ; Limpopo belt between Zimbabwe and Kaapvaal cratons Ring, 1993; Stuart and Zengini, 1987;. Ž . Ž .McCourt and Vearncombe, 1987; Droop, 1989 ; Namaqua belt along the south boundary of K Van der Merwe, 1995; Pearson, 1995 .

ated with times of mantle surge. The typical riftŽ .structure downdropped blocks of old rocks , thick

Ž .sequences of rift sediments Frostick et al., 1986and the narrow range of rift widths, typically 30–50km for the inner rift valley, are consistent withsub-rift downflow and an associated collapse mecha-nism of rift formation and cannot easily be adaptedto the popular Apull apartB hypothesis. A pull-apartdynamic system should be expressed as a variety ofrift widths, from narrow initial separations to maturerifts. No rift sequence of that kind is known.

Some segments of the rift system enclose upliftedhorsts on a scale that requires tectonic support. The

most striking example is the great uplifted mass ofRuwenzori, within the western rift. Snow-cappedRuwenzori, the second highest mountain in Africa, isa fault-bounded block that rises more than 4 km

Ž .above the surrounding plateau Holmes, 1965 , lead-ing one to imply tectonic support within a compres-sional regime. The western rift lies between the

ŽTanzanian and central African cratons Braile et al.,.1995, p. 219 , deduced sites of focused upwelling,

therefore is deduced to be a boundary where onemight expect convergent flow and downwelling.

The equatorial Tanzania craton is associated witha broad negative gravity anomaly, 1200 km wide,


amplitude 150"20 mGal, centered on Lake Victo-ria, and consistent with sub-craton upwelling. Theeastern and western rifts, around and between thecratons, are marked by narrow, steep-sided negative

Ž .anomalies Simiyu and Keller, 1996 . Their preferredw xgravity model op. cit., Figs. 8 and 9 , constrained by

seismic data, shows a deep mass of low densityasthenosphere below the craton keel, shallower low

Ždensity material beneath the rifts see below: seismic.data . That model is consistent with the regional heatŽ .flow Nyblade et al., 1990 , and with deflection of

upwelling flow around a thick cratonic keel, asŽ .proposed by Ballard and Pollack 1987 .

The Kenya rift, at the site of a recent seismicŽ .survey Green et al., 1991 is 50–70 km wide and

bounded by normal faults with offsets up to 4 km.The valley fill contains 2–3 km of volcanic andsedimentary materials. A seismic survey along arift-normal transect identified a narrow, steep-sidedlow-velocity zone in the sub-rift mantle. The investi-gators emphasized a Aremarkable correlationB be-tween the rift-bounding faults and the width of theupper mantle low-velocity zone in the depth range

Ž65–105 km. A later seismic survey Braile et al.,.1995 shows the narrow sub-rift LVZ extending to at

least 150 km depth. Green et al. attributed the LVZto decompression melting in a mantle plume, but itsnarrow dimensions and its location, directly belowthe rift, can more easily be accounted for by theproposed sub-rift sinking of low-density valley fillmaterial.

A recent preliminary investigation of steep-pathSKS waves traversing the region beneath the EastAfrican rifts showed that the polarization direction of

Žthe fast S-wave is parallel to the rift Ben-ismail and.Barruol, 1997 , an orientation that is inconsistent

with a Apull-apartB model of rift formation but is thatexpected for rift-normal uniaxial compression at asite of convergent flow. Further measurements ofSKS splitting in the region of the East African riftand of the Baikal rift, Siberia, showed a 908 changein the direction of polarization of the fast wave, fromnearly rift-normal in bordering regions to rift-parallel

Žwithin and near the rift Gao et al., 1997; Vauchez et.al., 1999 . The authors proposed rift-normal mantle

flow for the results in bordering regions, and Gao etŽ .al. 1997 attempted to account for rift-parallel polar-

ization of the fast split wave by the presence of

Ž .melt-filled microcracks. Vauchez et al. 1999 re-viewed the data and suggested rift-parallel mantleflow. In the light of the proposed alternative riftmodel, i.e. that rifts are collapse structures in regionsof convergence and downflow, the two sets of polar-ization directions, 908 apart, can be interpreted as theresult of mantle rheology at a flow corner, whereconvergent shear flow in bordering regions givesway to uniaxial compression and then to sub-riftdownflow. The olivine orientation in the near-riftcompressional regime is similar to that for a moun-

Ž . Ž .tain range cf. Fig. 23 , i.e. with the b axis sloworiented parallel to the direction of shortening, the aand c axes in a rift-parallel plane. That orientationaccounts for the rift-parallel fast S-wave and obvi-ates the need to appeal to rift-parallel mantle flow.

3.3. Sub-rift recycling

3.3.1. DiÕersity of rift ÕolcanismSupporting evidence of sub-rift downflow is pro-

vided by geochemical indications that the composi-tion and variability of rift-related igneous rocks arerelated to sub-rift recycling of a variable mixture ofmantle-source material with crust and rift valleysediments. Interpretations are faced with uncertaintydue to the similar seismic anomalies from zones ofpartial melting and zones of hydration in the uppermantle. A recent seismic survey of the Afar region,the connecting link between the oceanic and conti-nental rift systems, delineated an extensive LVZ withP-wave retardations of ;2% and S-wave retardation

Ž .of ;4% Knox et al., 1998 . The authors attributedthe low velocities to elevated temperature and partial

Ž .melting, but the experiments of Christensen 1966show similar retardations for mantle peridotite with

Ž .about 11% of serpentine Section 1.1.12 .A significant clue to rift dynamics is the abrupt

shift in igneous rock diversity at the ocean to conti-nent boundary. In contrast with the relative uniform-ity of MORB, the volcanic and intrusive rocks asso-ciated with continental rifts include not onlyMORB-like and alkalic basalts but a wide range of

Žother igneous rocks granite and rhyolite, for exam-ple, and various peralkalic silicate rocks with associ-

. Ž .ated carbonatites Kampunzu and Lubala, 1991 .The observed diversity is difficult to account for interms of sub-axial upwelling and decompressionmelting of upper mantle peridotite, because partial


melting, in general, tends to yield a uniform product,constrained by the composition of the thermal mini-mum in the system. The diversity of rift-relatedigneous rocks can more easily be accounted for if wespecify a source within heterogeneous, Amarble cakeBmantle that developed by multiple recycling of crustalmaterials. In accord with that model much of themagmatic diversity can be attributed to recycling ontwo scales: long-range recycling of oceanic crust andintercalated sediments via major return flow gyres,and short-range recycling, via secondary sub-conti-nent convection, of continental materials, includingterrigenous detritals, earlier-formed extrusives, plusevaporites and chemical precipitates deposited in riftvalleys. The broad question of rift-related igneousdiversity is complex, treated in an extensive litera-ture, and beyond the limits of the present paper. Thepresent review is restricted to Africa and is focusedon the carbonatite:alkali silicate association that is acharacteristic feature of continental rifts and mobilebelts but is notably absent at mid-ocean ridges andvolcanic arcs.

3.3.2. Helium isotopic ÕariationThe increased petrographic diversity associated

with the ocean to continent rift extension is matchedby increased variability of the helium isotope ratios

Ž .of basalt Scarsi and Craig, 1996 . Oceanic basalts ofthe Gulf of Aden have the uniform isotope ratiosŽ .RrR s8"1 that are typical of MORB in allA

oceans. In notable contrast, the helium isotope ratiosof basalts in the Afar Depression and Ethiopian Rift,continental extensions of the mid-ocean rift, the he-lium isotope ratios are widely variable, from 5 to 17

Ž y6 .times the atmospheric value R s1.4=10 .A

Scarsi and Craig attributed the high ratios to contri-butions from a deep-source plume of APrimitiveMantleB enriched in primordial helium-3. A reason-able alternative, consistent with the proposed sub-riftsubduction and recycling, is that the high heliumisotope ratios are the signature of sub-aerial basaltenriched in cosmogenic helium-3 and that the lowratios reflect an increased contribution from conti-nental material enriched in radiogenic helium-4.

3.3.3. Genesis of rift carbonatitesDetailed descriptions and various genetic hy-

potheses for carbonatites are thoroughly covered bya collection of papers in the carbonatite volume

Ž .edited by Bell 1989 . The carbonatite:alkalic silicateoccurrences can be referred to as carbonatites forpurposes of discussion although the typical showingshave a much larger volume of the alkalic silicaterocks: nepheline syenite and ijolite intrusives, forexample and volcanic equivalents such as phonolite.The igneous carbonate rocks typically are emplacedafter the more voluminous alkali silicate rocksŽ .Barker, in Bell, 1989 .

Most investigators of carbonatites, influenced bydistinctive trace element and isotopic signatures, fa-vor a mantle origin. The carbon isotopic compositionof carbonatites has a wide range, y2.4 to y7.9‰,and a mean value of y5.1‰, indistinguishable fromthat of kimberlites and diamonds, which have usually

Žbeen taken as samples of deep-seated carbon Deines.and Gold, 1973 . One model is based on fractional

crystallization of a carbonated alkalic magma, an-other on separation of an immiscible carbonate melt

Žfrom a carbonated peralkaline magma see papers in.Bell, 1989 . Neither model offers a convincing ex-

planation as to how the coexisting carbonate meltsand alkalic silicate melts can be produced by partialmelting of mantle peridotite, in which carbon is arare trace element. A proposed alternative, consistent

Žwith a general recycling model of volcanism Keith,.1993 , is that rift-related igneous rocks, including

carbonatites, are products of partial melting and reac-tions in a sub-rift zone of downflow and mixing. Thecarbonatites can be attributed to melting of sedimen-tary carbonate rocks and to reactions between car-bonate melts and other materials in the recycledmixture. Experimental fusions show that limestonesyield a melt fraction at temperatures in the range600–6758C, near the eutectic in the system CaO–

Ž .MgO–CO –H O Lentz, 1998 , and that some of2 2

the isotopic and trace element signatures are ac-quired during reactions between carbonate melt andvarious intruded rocks. The proposed genetic linkwith sedimentary carbonates is consistent with sev-eral lines of evidence. The major element features ofcarbonatites, although modified by reactions, are the

Žbasis for subdivision into three classes calcic, mag-.nesian and ferruginous that are clearly related to the

principal types of sedimentary carbonates and theirabundance ratios.

The proposed process is consistent with a generalŽ .recycling model of volcanism Keith, 1993 , and the


reactions are similar to those postulated for theAlimestone syntexisB model that was proposed someyears ago for the origin of peralkaline igneous rocksby desilication of granitic or basaltic magmas at

Žcontacts with intruded limestone Shand, 1930; Daly,.1933; Tilley, 1952 . The proposed recycling model

differs mainly in that reactions are proposed to occurin subduction zone mixtures, with resultant cogenera-tion of carbonatites and peralkaline silicate rocks. Itis presumed that the carbonate melt may be producedfrom remotely recycled marine limestone or fromlocally recycled lacustrine carbonates.

The classic examples of limestone syntexis are atsites where limestone or dolomite was intruded by

Ž .magma either granitic or basaltic , and where thereactions left a dual record: the formation of Ca, Mgsilicates in the contact zone of the limestone, alkalienrichment and desilication in the intrusive, andeventual development of feldspathoidal minerals. Itis a small step from that type of occurrence to theprocess that is here proposed to yield carbonatites, aprocess in which the carbonate reactant is fused andthe desilication reactions occur between carbonatemagma and silicate minerals or magma. In that casethe direct, in-situ evidence is lost, of course, and thegenetic process must be deduced from the geochem-istry of the products: carbonatites, alkalic silicate

Ž .rocks, alkalic alteration haloes AfenitesB andevolved gases, dominated by CO .2

The proposed recycling model for carbonatites istied to a general model of mantle convection inwhich upwelling of sub-continental mantle is fo-cused beneath the thick lithospheric keels of ancientPrecambrian cratons. Some parts of the heteroge-neous Amarble cakeB mantle, particularly thosemasses that contain recycled hydrous minerals orsedimentary carbonates, will be subject to volatile-promoted melting at the depths of dehydration ordecarbonation reactions. That process has been pro-posed to produce carbonate-rich magmas in the man-

Ž .tle Wyllie and Wolf, 1993 , and to favor theirponding at the base of the sub-cratonic lithosphereŽ .150–250 km depth .

Fig. 25 shows the locations of African carbon-atites and serves to bring out their association with

Ž .the borders of ancient cratons cf. Wooley, 1989 . Anorthern to mid-African group is clearly associatedwith rift systems, notably those surrounding the Tan-

zania craton, and a southern group mainly associatedwith mobile belts bordering or between other cra-tons. It is proposed that deflection of the residualsub-cratonic plume to shield borders and intermix-ture with rift valley sediments etc., in sub-rift sub-duction zones, will provide a potential source of awide variety of magmas, including carbonatites. The

Ž .proposed parental sediments are of two classes: 1marine limestone intercalated with oceanic crust atthe mid-ocean ridge axial zone and recycled, ocean

Ž .to continent, by major return-flow gyres, and 2carbonates precipitated as evaporite deposits in rift-valley lakes and recycled by local sub-rift downflow.It can be shown that carbonatites from these twoclasses of sedimentary source rocks can be differen-tiated, despite extreme compositional changes duringdecarbonation reactions.

There are indications, from the dominant CO2

content of MORB volatiles, that much of the interca-lated carbonate in oceanic crust is subject to decar-bonation reactions during slow recycling understeady-state conditions. During a surge of mantleflow, however, the sub-axial subduction column pre-sumably will be at lower temperatures and some ofthe recycled carbonate will escape decarbonation andwill be available for incorporation in the major re-turn flow gyre. Eventually it will be a component ofheterogeneous mantle upwelling beneath the ancientcontinental shields, followed by downwelling belowsurrounding rift systems and mobile belts. The pro-posed sedimentary carbonate recycling model is con-sistent with a recent study of the distribution of

Ž .carbonatites through time Veizer et al., 1992 . DatedŽ .occurrences of carbonatite Ns330 fall into groups

that appear to linked with major orogenic events, anassociation that fits in with the view that orogeniesare related to surges of mantle flow and increased

Žrecycling of crustal and sedimentary materials Keith,.1993 .

3.3.4. Carbon dioxide emissionsSome aspects of the genesis of rift-related carbon-

atites and alkalic volcanics can be related to recentdata regarding volcanic CO emissions. In compari-2

son with arc volcanoes, characterized by low CO2

volumes, MORB volcanism is presumed to be aŽ .major source Section 2.1.2 . Good measurements

are lacking but CO is the principal MORB volatile,2


a feature that can be attributed to dominant carbonatesediments interlayered with lavas at the axial sites ofrecycling. The high carbon dioxide emissions associ-ated with the alkalic volcanism of continental riftsystems can be attributed to the proposed formationof alkalic silicate magmas by desilication reactionsthat produce copious carbon dioxide. Typical mea-

Žsurements and estimates, in molryear with charac-. 9teristic petrology in parenthesis , include 520=10

Ž .molryear for Mt. Etna, Sicily alkalic basalt , 390=109 for East African rift volcano NyiragongoŽ . 9nephelinite , and 60=10 for Oldoinyo Lengai vol-

Ž .cano in the Eastern rift carbonatite and phonoliteŽ .Allard et al., 1991; Brantley and Koeppenick, 1995 .

In the light of the proposed recycling model ofvolcanism, the contrasts in CO emission can be2

related to the availability of carbonates in the recy-cling mix. The low CO emission from arc volca-2

noes, and their low CO rH O ratios are attributed to2 2

the carbonate deficiency of the recycled sediments:the deep-trench sediments that accumulate well be-low calcite compensation depth. At the other ex-treme, the high CO emission from continental rift2

volcanoes is attributed to sub-rift decarbonation reac-tions involving fused masses of sedimentary carbon-

Ž .ates from local sources lacustrine as well as remoteŽ .sources marine carbonate .

3.3.5. Oldoinyo Lengai, key to the rift systemKey evidence for a recycling origin of carbon-

atites is provided by the unusual carbonatite lavas ofOldoinyo Lengai volcano in the eastern rift, the onlyknown example of an active carbonate volcano. TheLengai lavas include both alkalic carbonate flowsand alkalic silicate flows, apparently immiscibleŽ .Dawson, 1989; Dawson et al., 1992 . The carbonateflows are unique in that they are dominantly sodic, incontrast with the common varieties of intrusive car-bonatite, dominantly calcic, and their extrusion tem-

Ž .peratures. 573–5938C and viscosities are extremelyŽ .low, the latter as low as 105 Pa s Dawson, 1989 .

The composition of Lengai lavas is near the sodicŽend of the system Na CO –K CO –CaCO Cooper2 3 2 3 3

.et al., 1975 . Principal mineral phases are Nyererite,Ž .Na Ca CO , and a potassium analogue, both of2 3 2

which melt incongruently at temperatures near 8008C.The liquidus surface in the ternary system has aminimum at 6658.

Most of the published genetic hypotheses for theLengai lavas have led to the conclusion that they aretrue carbonatites, despite their unusual composition.Their isotopic and trace element signatures, signifi-cantly different from sedimentary carbonates, havebeen interpreted as indicating a mantle origin for the

Ž .parental melt Bell and Keller, 1995 . The obviousdifficulty of generating Lengai-type carbonatite bypartial melting of mantle peridotite rekindles interestin an alternative hypothesis, proposed by MiltonŽ .1968, 1989 , that the Lengai natrocarbonatite lavasare not true igneous carbonatites but are the result ofnephelinitic magma intruding and mobilizing the lo-cal lacustrine carbonates. Evaporite beds, mainlycomposed of trona, Na CO PNaHCO P2H O and2 3 3 2

interbedded volcanic ash, occur in nearby rift valleylakes Natron and Magadi, and trona is producedcommercially from deposits up to 40 m thick that

2 Žcover 75 km in the Lake Magadi basin Jones et al.,.1977 . The Milton hypothesis takes on accentuated

significance in the light of the proposed generalmodel of sub-rift downflow and recycling, and isconsistent with the evidence of short-lived uraniumseries disequilibrium in the Lengai lavas, taken to

Ž .indicate a recent origin Pyle et al., 1991 .The proposed hypothesis, utilizing the Oldoinyo

Lengai geology as prime evidence, can be consideredas a modification of the Milton hypothesis, but dif-fers in that it is postulated that the Oldoinyo lavasare true carbonatites, though unusual, and that allcarbonatites and associated feldspathoidal silicate

Ž .rocks nephelinite at Oldoinyo Lengai can be at-tributed to reactions involving fused sedimentarycarbonates and a variety of siliceous rocks and mag-mas. The unusual compositional features of theLengai lavas can be attributed to a different primarysource of carbonate flux: lacustrine trona for Lengaicarbonatite, recycled marine limestone for mostcarbonatites. The commonly associated peralkalinesilicate rocks, typically containing feldspathoidalminerals such as nepheline, can be attributed todesilication reactions between the carbonate flux andsilicate rocks or silicate magma. Lengai lavas pro-vide a key to understanding that process becausethey are relatively young, because they represent anintermediate stage of those reactions, and becausetheir unique sodic character confirms their geneticassociation with the local sodic evaporite sediments.


Available sediments in the Natron and Magadi riftvalley basins are of three classes. An older series ofvolcaniclastics, at least 50 m thick, deposited inancestral Lake Olorongo, include an intermediateseries of tuffs and clays, and a younger series ofevaporites that reach a thickness of at least 40 m in

Ž .Magadi Basin Eugster, 1986 . The Olorongo bedsconsist of zeolite-bearing volcaniclastic sands, siltsand clays interbedded with green cherts derived fromprecipitated magadiite, a hydrous sodium silicateŽ .Baker, in Frostick et al., 1986 . It has been sug-

Ž .gested Eugster, 1986, p. 177 , from the distributionof Olorongo sediments, that ancestral Lake Olorongowas a single large lake that probably filled both the

Natron and Magadi basins. The younger sedimentaryformations were deposited in separate restrictedbasins. It seems evident that there is a large presentlyavailable source of the sodic lacustrine sediments,and that they were even more abundant in recent pasttimes.

ŽPast investigators e.g., Suwa et al., 1975; Bell.and Keller, 1995 favored a mantle origin for carbon-

atites, including Oldoinyo Lengai, and rejected thetrona-source model of Milton, on the grounds thatthe Lengai trace element and isotopic signaturesshow similarities to more AnormalB carbonatites.There are indications, however, that the major com-positional features of carbonatites, including those

Ž .Fig. 26. Carbon and oxygen isotopic composition of the natrocarbonatite lava from Oldoinyo Lengai larger circles and a proposed sourceŽ . Ž .material, trona, Na CO NaHCO P2H O from rift valley evaporite beds in nearby Lake Magadi, southern Kenya smaller circles . Data2 3 3 2

Ž . Ž . Ž . Ž .from O’Neil and Hay 1973 open circles and Suwa et al. 1975 filled circles .Heavy-line arrowsproposed model of the trend of isotopic change due to decarbonation reactions: the total isotopic change across a contact

Ž . Ž .zone, from a marine limestone ', M , through contact zone marble to calcite crystallized within the intrusive mass ' , the Mount RoyalŽ .pluton of Quebec Deines and Gold, 1969 .

Dotted-line boxs the AOka boxB, which encompasses the typical isotopic compositions of the Oka carbonatite-alkalic intrusive of Quebec,Ž .as well as the typical low-end isotopic compositions of many other carbonatite complexes Deines, 1970 .

Ž .Dashed line, RsRayleigh fractionation line for similtaneous separation of two phases from a common reservoir Deines, 1989, p. 338 ,here proposed to represent residual carbonatite magma produced by the loss of isotopically heavy CO from a sedimentary carbonate source2

Ž .material see text .


from Lengai, can be attributed to the effective flux-ing action of carbonate melts, and to acquisition oftheir isotopic and trace element signatures duringreactions of carbonate melts with silicate materials.Plutonic xenoliths in the carbonatite are dominantly

Ž .alkalic nephelinites , with widely variable modes,typically show complex zoning and metasomatism,and mineral disequilibrium, attributed to derivation

Žwithin a CO -rich sub-volcanic complex Dawson et2.al., 1995 . Hypotheses based on derivation from

sedimentary carbonate source rocks have generallybeen rejected on the basis of significant differences

Ž .in isotopic ratios including Sr and in a selected listof characteristic trace elements, the Aindex elementsB

Ž .of Gold 1966 . The diagnostic significance of traceelement and isotopic features is open to question,however, on the grounds that many of the character-istic, AfingerprintsB of carbonatites are not primary.The strontium isotope ratios are subject to an addi-tional uncertainty because of the wide age-related

Ž .variation of the ratios in limestones Veizer, 1989 ,and because of considerable overlap between the

Žratios of carbonatites and limestones see Havatsu et.al., 1965, re Precambrian examples .

One of the principal composition features thatprior investigators have used as a basis for rejectingMilton’s trona-source hypothesis for Lengai carbon-

Ž .atites e.g. see Cooper et al., 1975 is that the carbonand oxygen isotope ratios do not match: the carbon-atites have appreciably lighter carbon and oxygen. It

Ž .is shown, however Fig. 26 , that the differences arein accord with expected changes due to decarbona-tion reactions and loss of CO . They follow a trend2

parallel to that of Rayleigh fractionation for simulta-Ž .neous separation of two phases Deines, 1989 , and

have magnitudes close to the total isotopic changeduring desilication reactions at an intrusive limestone

Ž .contact in Quebec Deines and Gold, 1969 . InŽthe Quebec example, the isotopic changes marine

limestone wall rock to residual calcite in the intru-. 13sive were d C . . . q1.1‰ to y4.9‰, and

d18 O . . . 24.1‰ to 10.1‰.

The trace element signatures are less conclusivein establishing a genetic relation between Lengaicarbonatite and the local trona beds, but it seems

Ž .clear Fig. 27 that the trace elements in the Lengaicarbonatite are those that can most easily be acquiredby the dissolution effects of a carbonate flux, and

Ž . Ž .Fig. 27. Trace element abundances in Lengai natrocarbonatite open circles and associated nephelinite filled triangles , normalized toŽ . Ž .primitive mantle concentrations, after Keller and Spettel, in Bell and Keller, 1995 , to show that the trace elements Th, Ta, Zr, and Ti that

are notably depleted in the Lengai natrocarbonate lava are from source-rock minerals that resist dissolution in sodium carbonate fluxŽ .Dolezal et al., 1968 .


conversely, that the notably depleted trace elementsŽ .Th, Ta, Zr and Ti; Fig. 27 are those that normallyoccur in minerals such as ilmenite, titanite and zir-con, which are relatively resistant to dissolutionŽ .Dolezal et al., 1968 .

4. Conclusions

It is concluded, based on rift continuity fromocean to continent, on rift valley structures, on geo-physical data and on the geochemistry of rift-relatedigneous rocks, that oceanic and continental rifts arecollapse structures, typically developed at zones ofconvergence and downflow in underlying mantle,and characterized by volatile-promoted selectiveremelting and reactions within a heterogeneous sub-rift mixture of mantle peridotite plus recycled sur-face materials. The process typically yields basalt inoceanic regions, and a variety of igneous rocks,including the key carbonatites, associated with conti-nental rifts. There is a wide array of structural,geophysical and geochemical evidence that is con-trary to the plate tectonics hypothesis. The abovethree-chapter summary is deliberately one sided, pre-sented with the intent of generating controversy, thelife blood of science. With similar intent I propose atopic for debate: resolÕed that plate tectonics is themost fantastic house of cards that has eÕer beenerected in science.

Endnotesa w xolivine crystallographic axis, pole to 100

face: propagation direction of fast P-wave,and vibration direction of fast S-wave

b w xolivine axis, pole to 010 face: propagationdirection of slowest P-wave

c w xolivine axis, pole to 001 faceflexload structure proposed term for a volcanic

geosyncline produced by differential load-ing

LNB level of neutral buoyancy in oceanic crust, alevel of balance between density of basalticmagma and the mean density of surround-ing upper crust

LNS Žlevel of neutral stress horizontal stresss.zero , the base of a Astress transitionB char-

acterized by crossovers of horizontal andvertical stress

Pn P-wave with a shallow path following thecurvature of the Earth

s , s , s Maximal and minimal horizontal stress,H h V

vertical stressV , VP S seismic velocities of P-wave, S-wave

Acknowledgements

The manuscript was improved as a result of criti-cal reviews of selected sections by Peter Deines,K.P. Furlong and E.K. Graham and discussions withother Penn State colleagues.

References

Abrams, L.J., Larson, R.L., Shipley, T.H., Lancelot, Y., 1993.Cretaceous volcanic sequences and volcanism. In: Pringle,

Ž .M.S., Sager, W.W., Sliter, W.V., Stein, S. Eds. , The Meso-zoic Pacific: Geology, Tectonics and Volcanism. AmericanGeophysical Union, Geophysical Monograph, vol. 77, pp.77–101.

Allard, P. et al., 1991. Eruptive and diffuse emissions of CO2

from Mount Etna. Nature 351, 387–391.Allegre, C.J., Staudacher, T., Sarda, P., Kurz, M.D., 1983. Con-

straints on evolution of Earth’s mantle from rare gas systemat-ics. Nature 303, 762–766.

Allegre, C.J., Staudacher, T., Sarda, P., 1986. Rare gas systemat-ics, formation of the atmosphere, evolution and structure ofthe earth. Earth and Planetary Science Letters 81, 127–150.

Allegre, C.J., Sarda, P., Staudacher, T., 1993. Speculations aboutthe cosmogenic origin of He and Ne in the interior of theearth. Earth and Planetary Science Letters 117, 229–233.

ŽAllen, R.M. et al., 1999. The Icelandic crust. Eos Transactions. Ž .American Geophysical Union 80 46 , F645.

Anderson, D.L., 1993. Helium-3 from the mantle: primordialsignal or cosmic dust? Science 261, 170–176.

Ž .Anderson, D.L., 1994. Plates and plumes. EOS 75 16 , 60.Anderson, D.L., 1998. The helium-heat-lead paradox. 7th Annual

V.M. Goldschmidt Conf., Tuczon, AZ, June 1997, vol. 921,Lunar and Planetary Science Inst., Houston, TX, p. 6.

Anderson, D.L., Tanimoto, T., Zhang, Y., 1992. Plate tectonicsand hotspots: the third dimension. Science 256, 1645–1651.

Armstrong, R.L., 1968. A model for the evolution of strontiumand lead isotopes in a dynamic Earth. Reviews of Geophysics6, 175–199.

Armstrong, R.L., 1991. The persistent myth of crustal growth.Australian Journal of Earth Sciences 38, 613–630.

Armstrong, R.L., Hein, S.M., 1973. Computer simulation of Pb


and Sr isotopic evolution of the Earth’s crust and uppermantle. Geochimica et Cosmochimica Acta 37, 1–18.

Auffret, G.A. et al., 1991. Hydrothermal and volcaniclastic de-w xposits on top of a Mid-Atlantic Ridge seamount abs. . Eos

Ž . Ž .Transactions, American Geophysical Union 72 44 , 470,Supplement.

Auzende, J.M. et al., 1989. Direct observation of a section throughslow-spreading oceanic crust. Nature 337, 726–729.

Auzende, J.M. et al., 1994. Observation of sections of oceaniccrust and mantle outcropping on the southern wall of the Kane

Ž .FZ N. Atlantic . Terra Nova 6, 143–148.Ave Lallemant, H.G., Carter, N.L., 1970. Syntectonic recrystal-´

lization of olivine and modes of flow in the upper mantle.Geological Society of America Bulletin 81, 2203–2220.

Babcock, J.M., Harding, A.J., Kent, G.M., Orcutt, J.A., 1998. Anexamination of along-axis magma chamber width and crustalstructure on the East Pacific Rise. Journal of GeophysicalResearch 103, 30451–30467.

Babuska, V., Cara, M., 1991. Seismic Anisotropy in the Earth.Kluwer, 217 pp.

Babuska, V., Plomerova, J., Sileni, J., 1987. Structural model ofthe sub-crustal lithosphere in Central Europe. In: Fuchs, K.,

Ž .Froidevaux, C. Eds. , The Composition, Structure and Dy-namics of the Lithosphere–Asthenosphere System, AGU Geo-dynamics Ser. a16, pp. 239–251.

Babuska, V., Plomerova, J., Sileny, J., 1993. Models of seismicanisotropy in the deep continental lithosphere. Physics ofEarth and Planetary Interiors 78, 167–191.

Baker, E.T., Massoth, G.J., Feely, E.A., Cannon, G.A., 1998. Therise and fall of the CoAxial hydrothermal site. Journal ofGeophysical Research 103, 9791–9806.

Ballard, S., Pollack, H.N., 1987. Diversion of heat by ArcheanŽcratons. Earth and Planetary Science Letters 85, 253–264 cf.

.Journal of Geophysical Research, v. 92, p. 6291 .Ballard, R.D., Van Andel, Tj.H., 1977. Morphology and tectonics

of the inner rift valley on the Mid-Atlantic Ridge. GeologicalSociety of America Bulletin 88, 507–530.

Bamford, D., 1977. Pn velocity anisotropy in a continental uppermantle. Geophysical Journal, Royal Astronomical Society 49,29–48.

Barnes, C., Cochran, J.C., 1988. The geochemistry of uranium inmarine sediments. In: Guary, J.C., Gueguiniat, P., Pentreath,

Ž .R.J. Eds. , Radionuclides: A Tool for Oceanography, pp.162–170.

Barruol, G., Souriau, A., Vauchez, A., Diaz, J., Gallart, J., Tubia,J., Cuavas, J., 1998. Lithospheric anisotropy beneath the Pyre-nees from shear wave splitting. Journal of Geophysical Re-

Ž .search B 103, 30039–30053.Barth, G.A., 1994. Plate boundary geometry to Moho depths

within overlapping spreading centers of the East Pacific Rise.Earth and Planetary Science Letters 128, 99–112.

ŽBaudry, N., Kroenke, L.W., 1991. Intermediate wavelength 400–.600 km South Pacific geoidal undulations, their relationship

to linear volcanic chains. Earth and Planetary Science Letters102, 430–436.

Beattie, P., 1993. Uranium–thorium disequilibria and partitioningŽ .on melting of garnet peridotite. Nature London 363, 63–65.

Ž .Bell, K. Ed. , 1989. Carbonatites: Genesis and Evolution. UnwinHyman, London.

Ž .Bell, K., Keller, J. Eds. , 1995. Carbonatite Volcanism: OldoinyoLengai and the Petrogenesis of Natrocarbonatites. Springer-Verlag.

Ben-Ismail, W., Barruol, G., 1997. Upper mantle structure of theŽAfrican plate from SKS splitting. Eos Transactions, American

. Ž .Geophysical Union 78 46 , F486.Bennett, V.C., Esat, T.M., Norman, M.D., 1996. Two mantle–

plume components in Hawaiian picrites inferred from corre-lated Os–Pb isotopes. Nature 381, 221–223.

Bercovici, D.G, Schubert, G., Glatzmaier, G.A., 1989a. Influenceof heating mode on three-dimensional mantle convection.Geophysical Research Letters 16, 617–620.

Bercovici, D., Schubert, G., Glatzmaier, G.A., 1989b. Three-di-mensional spherical models of convection in the Earth’s man-tle. Science 244, 950–955.

Berner, R.A., Caldeira, K., 1997. The need for mass balance andfeedback in the geochemical carbon cycle. Geology 25, 955–

Ž .956 see comment by M. J. Bickle in v. 26, p. 477 .Bickle, M.J., 1996. Metamorphic decarbonation, silicate weather-

ing, and the long-term carbon cycle. Terra Nova 8, 270–276.Bina, C.R., Helffrich, G., 1994. Phase transition Clapeyron slopes

and transition zone seismic discontinuity topography. Journalof Geophysical Research 99, 15853–15860.

Blackman, D.K., Orcutt, J.A., Forsyth, D.W., Kendall, J.M., 1993.Seismic anisotropy in the mantle beneath an oceanic spreadingcenter. Nature 366, 675–677.

Blackman, D.K., Cann, J.R., Janssen, B., Smith, D.K., 1998.Origin of extensional core complexes: evidence from theMid-Atlantic Ridge at Atlantis fracture zone. Journal of Geo-physical Research 103, 21315–21333.

Blank, J.G., Delaney, J.R., DesMarais, D.J., 1993. The concentra-tion and isotopic composition of carbon in basaltic glassesfrom Juan de Fuca Ridge, Pacific Ocean. Geochimica etCosmochimica Acta 57, 875–887.

Bock, G., 1991. Long-period S to P converted waves and theonset of partial melting beneath Oahu, Hawaii. GeophysicalResearch Letters 18, 869–872.

Bodvarsson, G., Walker, G.P.L., 1964. Crustal drift in Iceland.Geophysical Journal of the Royal Astronomical Society 8,285–300.

Bohlen, S.R., Liu, J., Ernst, W.G., Liou, J.G., 1994. Stability ofhydrous phases and source of water in deep subduction zonesw x Ž .abs. . Eos Transactions, American Geophysical Union 75Ž .44 , 748, Supplement.

Bonatti, E., 1976. Serpentinite protrusions in the oceanic crust.Earth and Planetary Science Letters 32, 107–113.

Bonatti, E., 1990. Not so hot AhotspotsB in oceanic crust. Science250, 107–110.

Bonatti, E., Chermak, A., 1981. Formerly emerging crustal blocksin the equatorial Atlantic. Tectonophysics 72, 165–180.

Bonatti, E., Honnorez, J., 1971. Non-spreading crustal blocks atthe Mid-Atlantic Ridge. Science 174, 1329–1331.

Bonatti, E., Honnorez, J., 1976. Sections of the Earth’s crust inthe equatorial Atlantic. Journal of Geophysical Research 81,4104–4116.


Bonatti, E., Emiliani, C., Ferrara, G., Honnorez, H., Rydell, H.,1974. Ultramafic-carbonate breccias from the equatorial Mid-Atlantic Ridge. Marine Geology 16, 83–102.

Bonatti, E., Honnorez-Guerstein, M.B., Honnorez, J., Stern, C.,1976. Hydrothermal pyrite concretions from the Romanchefracture zone: Metallogenesis in oceanic fracture zones. Earthand Planetary Science Letters 32, 1–10.

Bonatti, E., Chermak, A., Honnorez, J., 1979. Tectonic andigneous emplacement of crust in ocean transform zones. In:

Ž .Talwani, M., Harrison, C.G., Hayes, D.E. Eds. , Deep DrillingResults in the Atlantic Ocean: Ocean Crust. Scientific Report48, American Geophysical Union, Washington, D.C, pp. 239–248.

Bonatti, E. et al., 1992. Romanche transformrMid-Atlantic Ridgew x Žeastern intersection abs. . Eos Transactions, American Geo-

. Ž .physical Union 73 14 , 295, Supplement.Bonatti, E., Ligi, M., Gasparini, L., Vera, E., 1994. Imaging

crustal uplift, emersion and subsidence at the Vema Fracturew x Ž .Zone abs. . Eos Translations, American Geophysical Union

Ž .75 32 , 371–372, Supplement.Bonatti, E., Ligi, M., Borsetti, A.M., Gasparini, L., Negri, A.,

Sartori, R., 1996. Lower Cretaceoous deposits trapped near theequatorial Mid-Atlantic Ridge. Nature 380, 518–520.

Booij, E., Staudigel, H., 1996. Bulk composition estimates ofw x Žmature oceanic crust abs. . Eos Transactions, American Geo-

. Ž .physical Union 77 46 , 781, Supplement.Borman, P., Burghardt, P.T., Makeyeva, L.I., Vinnick, L.P., 1993.

Teleseismic shear-wave splitting and deformations in CentralEurope. Physics of Earth and Planetary Interiors 78, 157–166.

Bose, K., Navrotsky, A., Schields, P.J., Weidner, D.J., 1996.Calorimetry, phase equilibria and equation of state mesure-ments of hydrous magnesium silicates and implications for

w x Žtransport to the mantle abs. . Eos Transactions, American.Geophysical Union 77, S281.

Bott, M.P., 1971. The Interior of the Earth. Edward Arnold,London.

Bottinga, Y., Javoy, M., 1989. MORB degassing: evolution ofCO . Earth and Planetary Science Letters 95, 215–225.2

Bottinga, Y., Steinmetz, L., 1979. A geophysical, geochemical,petrological model of the submarine lithosphere. Tectono-physics 55, 311–347.

Bowin, C., 1973. Origin of Ninetyeast Ridge from studies near theequator. Journal of Geophysical Research 78, 6029–6043.

Boyd, F.R., Gurney, J.J., 1986. Diamonds and the African litho-sphere. Science 232, 472–477.

Braile, L.W., Keller, G.R., Wendlandt, R.F., Morgan, P., Khan,M.A., 1995. The East African rift system. In: Olsen, K.H.Ž .Ed. , Continental Rifts: Evolution, Structure, Tectonics. Else-vier, Amsterdam.

Brantley, S.L., Koeppenick, K.W., 1995. Measured carbon diox-ide emissions from Oldoinyo Lengai and the skewed distribu-tion of passive volcanic fluxes. Geology 23, 933–936.

Brey, G., Brice, W.R., Ellis, D.D., Green, D.H., Harris, K.L.,Ryabchikov, I.D., 1983. Pyroxenecarbonate reactions in theupper mantle. Earth and Planetary Science Letters 62, 63–68.

Burnham, C.W., 1981. Convergence and mineralization—is there

a relation? Geological Society of America, Memoir 154, 761–768.

Burr, N.C., Solomon, S.C., 1978. The relationship of sourceparameters of oceanic transform earthquakes to plate velocityand transform length. Journal of Geophysical Research 83,1193–1205.

Busse, F.H., Whitehead, J.A., 1971. Instabilities of convection ina high Prandtl number fluid. Journal of Fluid Mechanics 47,305–320.

Cande, S.C., 1978. Anomalous behavior of the paleomagneticfield inferred from the skewness of anomalies 33 and 34. Earthand Planetary Science Letters 40, 275–286.

Cande, S.C., Kent, D.V., 1985. Tectonic rotations and theirpaleomagnetic consequences for oceanic basalts: comment.Journal of Geophysical Research 90, 4647–4651.

Cann, J. et al., 1992. Major landslides in the MAR median valley,w x Ž25–308N abs. . Eos Transactions, American Geophysical

. Ž .Union 73 43 , 569, Supplement.Cannat, M., 1993. Emplacement of mantle rocks in the seafloor of

Ž .mid-ocean ridges. Journal of Geophysical Research B 98,4163–4172.

Cannat, M., Mervel, C., Stakes, D., 1991. Stretching of the deepcrust at the slow-spreading Southwest Indian Ridge. Tectono-physics 190, 73–94.

Castillo, P.R., 1988. The Dupal anomaly as a trace of upwellinglower mantle. Nature 336, 667–670.

Castillo, P.R., Floyd, P.A., France-Lenord, C., 1992. Isotopegeochemistry of Leg 129 basalts: implications for the origin ofwidespread Cretaceous volcanic events in the Pacific. Proceed-ings of the Ocean Drilling Program, Scientific Results 129,405–413.

Castillo, P.R., Pringle, M.S., Carlson, R.W., 1994. East MarianaBasin tholeiites: cretaceous basalts related to the Ontong Javaplume? Earth and Planetary Science Letters 123, 139–154.

Castle, J.C., Creager, K.C., 1998. Topography of the 660-kmseismic discontinuity beneath Izu-bonin: Implications for tec-tonic history and slab deformation. Journal of Geophysical

Ž .Research B 103, 12511–12527.Cazenave, A., 1992. Geosat-derived geoid anomalies at medium

wavelength. Journal of Geophysical Research 97, 7081–7096.Cazenave, A., Monnereau, M., Gibert, D., 1987. Seasat gravity

undulations in the central Indian Ocean. Physics of the Earthand Planetary Interiors 48, 130–141.

Cerling, T.E., Craig, H., 1994. Geomorphology and in situ cosmo-genic isotopes. Annual Review of Earth and Planetary Sci-ences 22, 273–317.

Chaffey, D.J., Cliff, R.A., Wilson, B.M., 1989. Characterizationof the St. Helena magma source. In: Saunders, A.D., Norry, J.Ž .Eds. , Magmatism in the Ocean Basins, pp. 257–276, Geo-logical Society Special Publication no. 42.

Charlou, J.L. et al., 1997. High methane flux on the Mid-Atlanticw xRidge; hydrothermal and serpentinization processes abs. . Eos


Charlou, J.L. et al., 1998. Intense CH4 plumes generated byserpentinization of ultramafic rocks at the intersection of the


15820XN fracture zone and the Mid-Atlantic Ridge. Geochim-ica et Cosmochimica Acta 62, 2323–2333.

Chase, C.G., 1981. Oceanic island Pb: two-stage histories andmantle evolution. Earth and Planetary Science Letters 52,277–284.

Chauvel, C., Hofmann, A.W., Vidal, P., 1992. HIMU-EM: theFrench Polynesian connection. Earth and Planetary ScienceLetters 110, 99–119.

Chesnokov, Y.M., 1973. On elastic anisotropy of multicomponentmodels of the upper mantle structure. Physics of the SolidEarth 5, 290–296.

Chevrot, S., Vinnik, L., Montagner, J.-P., 1999. Global-scaleanalysis of the mantle Pds phases. Journal of GeophysicalResearch 104, 20203–20219.

Chopra, P.N., Paterson, M.S., 1984. The role of water in thedeformation of dunite. Journal of Geophysical Research 89,7861–7876.

Christensen, N.I., 1966. Elasticity of ultrabasic rocks. Journal ofGeophysical Research 71, 5921–5931.

Christensen, N.I., 1984. The magnitude, symmetry and origin ofmantle anisotropy based on fabric analysis of ultramafic tec-tonites. Geophysical Journal, Royal Astronomical Society 76,89–111.

Christensen, U.R., Harder, H., 1991. 3-D convection with variableviscosity. Geophysical Journal International 104, 213–226.

Christensen, N.I., Lundquist, S.M., 1982. Pyroxene orientationwithin the upper mantle. Geological Society of America,Bulletin 93, 279–288.

Christensen, U.R., Yuen, D.A., 1985. Layered convection inducedby phase transitions. Journal of Geophysical Research 90,10291–10300.

Clague, D.A., Dalrymple, G.B., 1988. Age and petrology ofalkalic post shield and rejuvenated stage lava from Kauai,Hawaii. Contributions to Mineralogy and Petrology 99, 202–218.

Clifford, T.N., 1966–1967. African structure and convection.Ž .Transactions of the Leeds Geological Society 7 5 , 291–301.

Cohen, R.R., Evensen, N.M., Hamilton, P.J., O’Nions, R.K.,1980. U–Pb, Sm–Nd and Rb–Sr systematics of mid-oceanridge basalt glasses. Nature 283, 149–153.

Cohen, A.S., O’Nions, R.K., Kurz, M.D., 1996. Chemical andisotopic variations in Mauna Loa tholeiites. Earth and Plane-tary Science Letters 143, 111–124.

Colliston, W.P., Prockelt, H.E., Schoch, A.E., 1991. A progres-sive ductile shear model of the Proterozoic Aggeneys Terrane,Namaqua mobile belt, South Africa. Precambrian Research 49,205–215.

Cooper, A.F., Gittins, J., Tuttle, O.F., 1975. The systemNa CO –K CO –CaCO at 1 kilobar, and its significance in2 3 2 3 3

carbonatite petrogenesis. American Journal of Science 275,534–560.

Cormier, M.H., Schierer, D.S., Macdonald, K.C., 1996. Evolutionof the East Pacific Rise and bisection of overlapping spreadingcenters by new, rapidly propagating ridge segments. MarineGeophysical Research 18, 53–84.

Courtney, R.C., White, R.S., 1986. Anomalous heat flow andgeoid across the Cape Verde Rise. Geophysical Journal of theRoyal Astronomical Society 87, 815–867.

Craig, H., Poreda, R.J., 1986. Cosmogenic 3He in terrestrialrocks: the summit lavas of Maui. Proceedings of the NationalAcademy of Sciences 83, 1970–1973.

Craig, H., Scarsi, P., 1998. High-3He hotspots: distribution andw xcorrelation with planetary scale mantle tomography abs . Eos

Ž . Ž .Transactions, American Geophysical Union 79 45 , 598,Fall Supplement.

Dahlen, F.A., 1982. Isostatic geoid anomalies on a sphere. Journalof Geophysical Research 87, 3943–3947.

Dalrymple, G.B., Lanphere, M.A., Clague, D.A., 1980. Conven-tional and 40Arr39Ar ages and the chronology of volcanicpropagation along the Hawaiian–Emperor chain. Initial Re-ports of the Deep Sea Drilling Program, vol. 55, U.S. Govern-ment Printing Office, Washington, DC, pp. 659–676.

Daly, R.A., 1933. Igneous Rocks and the Depths of the Earth.McGraw-Hill.

Davies, G.R., Norry, M.J., Gerlach, D.C., Cliff, R.A., 1989.Combined chemical and Pb–Sr–Nd isotopic study of theAzores and Cape Verde hot spots, in Magmatism in the Ocean

Ž .Basins. In: Saunders, A.D., Norry, M.J. Eds. , GeologicalSociety, Special Publication, vol. 42, pp. 231–255.

Davis, E.E., Lister, C.R., 1977. Tectonic structures on the Juan deFuca Ridge. Bulletin of the Geological Society of America 88,346–363.

Dawson, J.B., 1989. Sodium carbonate extrusions from OldoinyoLengai, Tanzania: implications for carbonatite complex gene-

Ž .sis. In: Bell, K. Ed. , Carbonatites: Genesis and Evolution.Unwin Hyman, London, pp. 255–277.

Dawson, J.B., Smith, J.V., Steele, I.M., 1992. The 1966 asheruption of the carbonatite volcano Oldoinyo Lengai: mineral-ogy of lapilli and mixing of silicate and carbonate magmas.Mineralogical Magazine 56I, 1–16.

Dawson, J.B., Smith, J.V., Steels, I.M., 1995. Petrology andmineral chemistry of plutonic igneous xenoliths from thecarbonatite volcano, Oldoinyo Lengai, Tanzania. Journal ofPetrology 36, 707–826.

Day, H.W., Bickford, M.E., 1989. The Nevadan orogeny in thenorthern Sierra Nevada: new constraints from U–Pb agesw x Ž .abs. . Eos Transactions, American Geophysical Union 70,490, Supplement.

Deines, P., 1970. The carbon and oxygen isotopic composition ofcarbonates from the Oka carbonatite, Quebec, Canada.Geochimica et Cosmochimica Acta 34, 1199–1225.

Deines, P., 1989. Stable isotope variations in carbonatites. In:Ž .Bell, K. Ed. , Carbonatites: Genesis and Evolution. Unwin

Hyman, London, pp. 301–359.Deines, P., Gold, D.P., 1969. The change in carbon and oxygen

isotopic composition during contact metamorphism of Trentonlimestone by the Mount Royal pluton. Geochimica et Cos-mochimica Acta 33, 421–424.

Deines, P., Gold, D.P., 1973. The isotopic composition of carbon-atites and kimberlite carbonates and their bearing on theisotopic composition of deep-seated carbon. Geochimica etCosmochimica Acta 37, 1709–1733.

Deplus, C. et al., 1998. Direct evidence of active deformation inthe eastern Indian Ocean plate. Geology 26, 131–134.

Derbyshire, F.A. et al., 1999. Crustal structure of central andnorthern Iceland from teleseismic receiver function analysis


w x Ž .abs. . Eos Transactions, American Geophysical Union 80Ž .46 , F645.

Detrick, R.S., White, R.S., Purdy, G.M., 1993. Crustal structure ofNorth Atlantic fracture zones. Reviews of Geophysics 31,439–458.

Dewey, J.F., Bird, J.M., 1970. Mountain belts and the new globaltectonics. Journal of Geophysical Research 75, 2625–2647.

Dewey, J.F., Pitman, W.C., Ryan, W.B.F., Bonnin, J., 1973. Platetectonics and the evolution of the Alpine system. GeologicalSociety of America, Bulletin 84, 3137–3180.

Dick, H.J. et al., 1991. Tectonic evolution of the Atlantis IIfracture zone, Indian Ocean. Proceedings of the Ocean DrillingProgram, Scientific Results, Leg 118, College Station, pp.359–398.

Doe, B.R., 1994. Zinc, copper and lead in mid-ocean ridgebasalts, and the source rock control on ZnrPb in ocean ridgehydrothermal deposits. Geochimica et Cosmochimica Acta 58,2215–2223.

Doin, M.P., Fleitout, L., Christensen, U., 1996. Stability of thicklithospheric roots in mantle convection: numerical experimentsw x Ž .abs. . Eos Transactions, American Geophysical Union 77Ž .46 , 751, Supplement.

Dolezal, J., Provondra, P., Sulcek, Z., 1968. Decomposition tech-niques. Inorganic Analysis. American Elsevier Publishing, pp.93–102.

Dong, H., Hall, C.M., Peacor, D.R., Halliday, A.N., 1995. Mecha-nisms of argon retention in clays revealed by laser 40Arr39Ardating. Science 267, 355–357.

Donnelly, T., Francheteau, J., Bryan, W., 1980. Initial Reports,Deep Sea Drilling Program, vol. 51, U.S. Government PrintingOffice, Washington, p. 75.

Douville, E., 1997. Trace elements in fluids from the new Rain-bow hydrothermal field: a comparison with other Mid-Atlantic

w x ŽRidge fluids abs. . Eos Transactions, American Geophysical. Ž . wUnion 78 46 , 832, Supplement cf. related abstracts on pp.

x831 .Droop, G.T., 1989. Reaction history of garnet-sapphirine gran-

ulites and conditions of high pressure metamorphism in thecentral Limpopo belt, Zimbabwe. Journal of MetamorphicPetrology 7, 383–403.

Duffield, W.A., 1978. Significance of contrasting vesicularity inbasalt from DSDP sites 407, 408 and 409, Reykjanes Ridge.

Ž .In: Luyendyk, B.P., Cann, J.R. Eds. , Initial Reports of theDeep Sea Drilling Project, vol. 49, U.S. Government PrintingOffice, Washington, pp. 715–719.

Duncan, R.A., Clague, D.A., 1985. Pacific plate motion recordedby linear volcanic chains. In: Nairn, A.E., Stehli, F.G., Uyeda,

Ž .S. Eds. , The Ocean Basins and Margins, vol. 7A, pp. 89–121.Duncan, R.A. et al., 1991. Age distriubtion of volcanism along

aseismic ridges in the eastern Indian Ocean. Proceedings ofthe Ocean Drilling Program, Scientific Results 121, 507–517.

Dupre, B., Allegre, C.J., 1983. Pb–Sr variation in Indian Oceanbasalts and mixing phenomena. Nature 303, 142–146.

Dziewonski, A.M., Forte, A.M., Su, W.-J., Woodward, R.L.,1993. Seismic tomography and geodynamics. In: Aki, K.,

Ž .Dmowska, R. Eds. , Relating Geophysical Structures andProcesses. American Geophysical Union, Geophysical Mono-graph, vol. 76, pp. 67–105.

Ebinger, C.J., Rosendahl, B.R., Reynolds, D.J., 1987. Tectonicmodel of the Malawi rift, Africa. Tectonophysics 141, 215–235.

Ebinger, C., Poudjom-Djomani, Y., Mbede, E., Foster, A., Daw-son, J.B., 1997. Rifting Archaean lithosphere: the Eyasi–Manyaro–Natron rifts, East Africa. Journal of the GeologicalSociety, London 154, 947–960.

Eggler, D.H., 1978. The effect of CO upon partial melting, with2

an analysis of melting in a peridotite–H O–CO system.2 2

American Journal of Science 278, 305–343.Eiler, J.M., Farley, K.A., Stolper, E.M., 1998. Correlated helium

and lead isotope variations in Hawaiian lavas. Geochimica etCosmochimica Acta 62, 1977–1984.

Einarsson, P., 1991. Earthquakes and present-day tectonism inIceland. Tectonophysics 189, 261–279.

Einarsson, P., Bransdottir, B., 1980. Seismological evidence forlateral magma intrusion during the 1978 deflation of Kraflavolcano in NE-Iceland. Journal of Geophysics 47, 160–165.

Eittreim, S.L., Gnibidenko, H., Helsey, C.E., Sliter, R., Mann, D.,Ragozin, N., 1994. Oceanic crustal thickness and seismiccharacter along a central Pacific transect. Journal of Geophysi-cal Research 99, 3139–3145.

Ekstrom, G., Dziewonski, A.M., 1998. The unique anisotropy inthe Pacific upper mantle. Nature 394, 168–172.

Elder, J., 1976. The Bowels of the Earth. Oxford Univ. Press.Eldholme, O., Grue, K., 1994. North Atlantic volcanic margins:

dimensions and production rates. Journal of Geophysical Re-search 99, 2955–2968 and 9263–9278.

Engelder, T., 1993. Stress Regimes in the Lithosphere. PrincetonUniv. Press.

England, P., Molnar, P., 1997. Active deformation of Asia: fromkinematics to dynamics. Science 278, 647–650.

Ernst, W.G., 1978. Petrochemical study of lherzolite rocks fromthe western Alps. Journal of Petrology 19, 341–392.

Ernst, W.G., 1999. Hornblende, the continent maker-evolution ofH O during circum-Pacific subduction. Geology 27, 675–678.2

Erzinger, J., Becker, K., Dick, H.J.B., Stokking, L.B., 1995.Proceedings, Ocean Drilling Program, Legs 137r140, pp.62–106, 293–312, College Station, TX.

Eugster, H.P., 1986. Lake Magadi, Kenya: a model for rift valleyhydrochemistry and sedimentation? In: Frostick, L.E., Renaut,

Ž .R.W., Reid, I., Tiercelin, J.-J. Eds. , Sedimentation in theAfrican Rifts. Geological Society Special Publication, vol. 25,pp. 177–189.

Evans, R.L. et al., 1998. The MELT experiment electromagneticw xdata; preliminary inversion and modeling results abs . Eos


Fagel, N., Andre, L., Debrabant, P., 1997. Multiple seawater-de-rived geochemical signatures in Indian Ocean pelagic clays.Geochimica et Cosmochimica Acta 61, 989–1008.

Farley, K.A., 1993. Is AprimordialB helium really extraterrestrial?Science 261, 166–169.

Farley, K.A., Craig, H., 1994. Atmospheric argon contaminationof oceanic island basalt olivine phenocrysts. Geochimica etCosmochimica Acta 58, 2509–2517.

Farley, K.A., Neroda, E., 1998. Noble gases in the Earth’s mantle.Annual Reviews of Earth and Planetary Sciences 26, 189–218.


Farley, K.A., Natland, J.H., Craig, H., 1992. Binary mixing ofŽ . Ženriched and undegassed primitive? Mantle components He,

.Sr, Nd, Pb in Samoan lavas. Earth and Planetary ScienceLetters 111, 183–199.

Feigenson, M.D., 1984. Geochemistry of Kauai volcanics and amixing model for the origin of Hawaiian alkali basalts. Contri-butions to Mineralogy and Petrology 87, 109–119.

Fisher, D.E., 1986. Rare gas abundances in MORB. Geochimicaet Cosmochimica Acta 50, 2531–2541.

Fisher, D.E., 1989. Evaluation of rare gas data in relation toŽ .oceanic margins. In: Saunders, A.D., Norry, M.J. Eds. , Mag-

matism in the Ocean Basins. Geological Society Special Publi-cation 42, pp. 301–311.

Fisher, D.E., 1994a. Mantle and atmospheric-like argon in vesi-cles of MORB glass. Earth and Planetary Science Letters 123,199–204.

Fisher, D.E., 1994b. Mantle Ar and Xe in MORB glass vesiclesw x Ž .abs. . Eos Transactions, American Geophysical Union 75Ž .16 , 233, Supplement.

Flanagan, M.P., Shearer, P.M., 1998. Topography on the 410-kmseismic velocity discontinuity near subduction zones fromstacking of sS, sP and pP precursors. Journal of Geophysical

Ž .Research B 103, 21165–21183.Ž .Fleming, T.H., Elliot, D.H., Foland, K.A., 1992. Isotopic Nd, Sr

and chemical characteristics of Ferrar tholeiites, Antarcticaw x Ž .abs. . Eos Transactions, American Geophysical Union 73Ž .14 , 329, Supplement.

Forsyth, D.W., 1992. Geophysical constraints on mantle and meltgeneration beneath mid-ocean ridges. In: Morgan, J.P., Black-

Ž .man, D.K., Swinton, J.M. Eds. , Mantle Flow and MeltGeneration at Mid-Ocean Ridges. Geophysical Monograph,vol. 71. American Geophysical Union, Washington, D.C., pp.1–65.

Forsyth, D.W., Webb, S.C., Dorman, L.N., Shen, Y., 1998. Phasevelocities of Rayleigh-waves in the MELT experiment, East

ŽPacific Rise. Science 280, 1235–1238 cf. related papers, p..1215 and 1221 .

Fox, P.J., Gallo, D.G., 1986. The geology of North Atlantictransform plate boundaries. In: Vogt, P.R., Tucholke, B.E.Ž .Eds. , The Western North Atlantic Region. Geological Soci-ety of America, Geology of North America, vol. M, pp.157–172.

Fox, P.J., Schreiber, E., Peterson, J.J., 1973. The geology of theocean crust: compressional wave velocities of oceanic rocks.Journal of Geophysical Research 78, 5155–5172.

Fox, P.J., Schreiber, E., Rowlett, H., McCamy, A., 1976. Thegeology of the Oceanographer fracture zone. Journal of Geo-physical Research 81, 4117–4128.

Frey, F.A., Garcia, M.O., Roden, M.F., 1994. Geochemical char-acteristics of Koolau volcano: implications of intershield geo-chemical differences among Hawaiian volcanoes. Geochimicaet Cosmochimica Acta 58, 1441–1462.

Frostick, L.E., Renaut, R.W., Reid, I., Tiercelin, J.J., 1986. Sedi-mentation in the African Rifts. Geological Society SpecialPublication, 25.

Fujimoto, H., Seama, N., Lin, J., Matsumoto, T., Tanaka, T.,

Fujioka, K., 1996. Gravity anomalies of the Mid-AtlanticRidge north of the Kane fracture zone. Geophysical ResearchLetters 23, 3431–3434.

Gallo, D.G., Fox, P.J., Macdonald, K.C., 1986. A sea beaminvestigation of the Clipperton fracture zone. Journal of Geo-physical Research 91, 3455–3467.

Gao, S. et al., 1997. SKS splitting beneath continental rift zones.Journal of Geophysical Research 102, 22781–22787.

Gary, M., McAfee, R., Wolf, C.L., 1965. Glossary of Geology.American Geological Institute, Washington, DC, 805 pp.

Gasperini, L. et al., 1997. Stratigraphic modelling of a carbonateplatform on the Romanche transverse ridge, equatorial At-lantic. Marine Geology 136, 245–257.

Gerlach, D.C., Cliff, R.A., Davies, G.R., Norry, M., Hodgson, N.,1988. Magma sources of the Cape Verdes archipelago: iso-topic and trace element constraints. Geochimica et Cos-mochimica Acta 52, 2979–2992.

Gladczenko, T.P., Hinz, K., Eldholme, O., Meyer, H., Neben, S.,Skogseid, J., 1997. South Atlantic volcanic margins. Journalof the Geological Society, London 154, 465–470.

Gold, D.P., 1966. The average and typical chemical compositionof carbonatites. Mineralogical Society of India, InternationalMineral Association Papers, 4th General, pp. 83–91.

Goldstein, S.J., Perfit, M.R., Batista, R., Murrell, D.J., Michael,T., 1994. Off-axis volcanism at the East Pacific Rise detectedby uranium series dating of basalts. Nature 367, 157–159.

Graham, D.W., Zindler, A., Kurz, M.D., Jenkins, W.J., Batista,R., Staudigel, H., 1988. He, Pb, Sr and Nd isotope constraintson mantle heterogeneity between young Pacific seamounts.Contributions to Mineralogy and Petrology 99, 446–463.

Graham, D.W., Christie, D.M., Harpp, K.S., Lupton, J.E., 1993.Mantle plume helium in submarine basalts from the Galapagosplatform. Science 262, 2023–2025.

Grand, S.P., Van der Hilst, R.D., Widiyantoro, S., 1997. Globalseismic tomography: a snapshot of convection in the Earth.

Ž .GSA Today 7 4 , 1–7.ŽGreen, H.W., 1995. Response to Bowen Award. Eos Transac-

. Ž .tions, American Geophysical Union 76 16 , 163–166.Green, W.V., Achauer, U., Meyer, R.P., 1991. A three-dimen-

sional seismic image of the crust and upper mantle beneath theKenya rift. Nature 354, 200–203.

Griggs, D., 1939. A theory of mountain building. AmericanJournal of Science 237, 611–650.

Gudmundsson, A., 1995. Stress fields associated with oceanictransform faults. Earth and Planetary Science Letters 136,603–614.

Hacker, B.R., 1995. Rate of blueschist and eclogite formation andthe rheology, buoyancy and seismicity of subducting oceanic

w x Ž .crust abs. . Eos Transactions, American Geophysical UnionŽ .76 46 , 535, Supplement.

Hager, B.H., Clayton, R.W., Richards, M.A., Comer, R.P.,Dziewonski, A.M., 1985. Lower mantle heterogeneity, dy-namic topography and the geoid. Nature 313, 541–545.

Haggerty, S.E., 1986. Diamond genesis in a multiply constrainedmodel. Nature 320, 34–38.

Haggerty, S.E., 1997. Diamond and high pressure rocks: clues to


w x Žthe geodynamics of Earth’s deep interior abs. . Eos Transac-. Ž .tions, American Geophysical Union 78 17 , 330, Supple-

ment.Haimson, B.C., 1979. New stress measurements in Iceland rein-

force previous hydrofracturing results: sH is perpendicu-maxw x Žlar to the axial rift zone abs . Eos Transactions, American

.Geophysical Union 60, 377.Haimson, B.C., Rummel, F., 1982. Hydrofracturing stress mea-

surements in the drill hole at Reydarfjordur, Iceland. Journalof Geophysical Research 87, 6631–6649.

Haimson, B.C., Voight, B., 1977. Crustal stress in Iceland. Pureand Applied Geophysics 115, 153–190.

Hall, J.M., Robinson, P.T., 1979. Deep crustal drilling in theNorth Atlantic Ocean. Science 204, 573–586.

Hamelin, B., Grousset, F., Sholkovitz, E.R., 1990. Pb isotopes insurficial sediments from the North Atlantic. Geochimica etCosmochimica Acta 54, 37–47.

Hamilton, E.L., 1956. Sunken Islands of the Mid-Pacific Moun-tains. Geological Society of America, Memoir, 64.

Hamilton, W.B., 1988. Plate tectonics and island arcs. Bulletin ofthe Eological Society of America 100, 1503–1527.

Hanan, B., Graham, D.W., 1996. Lead and helium isotope evi-dence from oceanic basalts for a common deep source ofmantle plumes. Science 272, 991–995.

Handy, M.R., 1990. The solid-state flow of polymineralic rocks.Journal of Geophysical Research 95, 8647–8661.

Hanyu, T., Kaneoka, I., 1997. Helium isotopic study of thew x ŽPolynesian HIMU abs. . Eos Transactions, American Geo-

. Ž .physical Union 78 46 , 838, Supplement.Harland, W.B., Armstrong, R.L., Cox, A.V., Craig, L.E., Smith,

A.G., Smith, D.G., 1990. A Geologic Time Scale 1989.Cambridge Univ. Press, p. 142.

Hart, S.R., 1984. A large isotope anomaly in the southern hemi-sphere mantle. Nature 309, 753–757.

Hart, S.R., 1988. Heterogeneous mantle domains: signatures, gen-esis and mixing chronologies. Earth and Planetary ScienceLetters 90, 273–296.

Hart, S.R., Hauri, E.H., Oschmann, L.A., Whitehead, J.A., 1992.Mantle plumes and entrainment: isotopic evidence. Science256, 517–520.

Hassanipak, A.A., Wampler, J.M., 1996. Radiogenic argon re-leased by stepwise heating of glauconite and illite: the influ-ence of composition and particle size. Clays and Clay Miner-als 44, 717–726.

Hast, N., 1973. Global measurements of absolute stress. Philo-sophical Transactions of the Royal Society of London 274A,

Ž .409–419 see data tabulated in Haimson and Voight, 1977 .Hauri, E.H., 1996. Variability in the Hawaiian mantle plume.

Nature 382, 415–419.Hauri, E.H., Hart, S.R., 1993. Re–Os systematics of HIMU and

EM2 oceanic island basalts from the S Pacific Ocean. Earthand Planetary Science Letters 114, 353–371.

Hauri, E.H., Hart, S.R., 1997. Rhenium abundances and systemat-ics in oceanic basalts. Chemical Geology 139, 185–205.

Hauri, E.H., Lassiter, J.C., DePaulo, D.J., 1996. Osmium isotopesystematics of drilled lavas from Mauna Loa, Hawaii. Journalof Geophysical Research 101, 11793–11806.

Havatsu, A., York, D., Farquhar, R.M., 1965. Significance of

strontium isotope ratios in theories of carbonatite genesis.Nature 207, 625–626.

ŽHaxby, W.F., 1985. Gravity field of the world’s ceans from.SEASAT radar altimetry . Lamont–Doherty Geological Ob-

servatory and U.S. Navy, Office of Naval Research, 1 sheet.Haxby, W.F., Weissel, J.K., 1986. Evidence for small scale

mantle convection from SEASAT altimeter data. Journal ofGeophysical Research 91, 3507–3520.

Head, J.W., Wilson, L., Smith, D.K., 1996. Mid-ocean ridgeeruptive events: evidence for dike widths, eruption rates andthe evolution of axial volcanic ridges. Journal of GeophysicalResearch 101, 28265–28280.

Heirtzler, J.R., Frawley, J.J., 1994. New gravity model for EarthŽ .science studies. GSA Today 4 11 , 269–270.

Hess, H.H., 1964. Seismic anisotropy of the uppermost mantle inthe Pacific. Nature 203, 629–631.

Hess, P.C., 1993. Phase equilibria constraints on the origin ofŽ .ocean floor basalts. In: Morgan, J.P. Ed. , Mantle flow and

Melt Generation at Mid-ocean Ridges. American GeophysicalUnion, Geophysical Monograph, vol. 71, pp. 67–102.

Hilton, D.R., Craig, H. et al., 1989. A helium isotope transectalong the Indonesian Archipelago. Nature 342, 906–908.

Hinz, K., 1981. Wedges of very thick oceanward-dipping layersbeneath passive continental margins. Geologische Jahrbuch, E22, 3–28.

Hirose, K., Kawamato, T., 1995. Hydrous partial melting oflherzolite at 1 GPa: the effect of H O on the genesis of2

basaltic magmas. Earth and Planetary Science Letters 133,463–473.

Hirth, G., Kohlstedt, D.L., 1996. Water in the oceanic uppermantle: implications for rheology, melt extraction and evolu-tion of the lithosphere. Earth and Planetary Science Letters144, 93–108.

Hiyagon, H., Ozima, M., 1986. Partition of noble gases betweenolivine and basalt melt. Geochimica et Cosmochimica Acta50, 2045–2057.

Hiyagon, H., Ozima, M., Marty, B., Zashu, S., Sakai, H., 1992.Noble gases in submarine glasses from mid-ocean ridges andLoihi seamount: constraints on the early history of the Earth.Geochimica et Cosmochmica Acta 56, 1301–1316.

Hoernle, K., Schmincke, H.-U., 1993. The role of partial meltingin the 15-Ma geochemical evolution of Gran Canaria: a blobmodel for the Canary hotspot. Journal of Petrology 34, 599–626.

Hofmann, A.W., 1997. Mantle geochemistry: the message fromoceanic volcanism. Nature 385, 219–229.

Hofmann, A.W., White, W.M., 1982. Mantle plumes from ancientoceanic crust. Earth and Planetary Science Letters 57, 421–436.

Holden, J.C., Vogt, P.R., 1977. Graphic solutions to problems ofw x Žplumacy abs. . Eos Transactions, American Geophysical

.Union 58, 573–580.Holmes, A., 1965. Principles of Physical Geology. The Ronald

Press, New York.Honda, M., McDougall, I., Patterson, D.B., Doulgeris, A., Clague,

A., 1991. Possible solar noble gas component in the Hawaiianbasalts. Nature 349, 149–151.

Honda, M., McDougall, I., Patterson, D.B., Doulgeris, A., Clague,


D.A., 1993a. Noble gases in submarine pillow basalt glassesfrom Loihi and Kilauea, Hawaii: a solar component in theEarth. Geochimica et Cosmochimica Acta 57, 859–874.

Honda, M., McDougall, I., Patterson, D., 1993b. Solar noble gasesin the Earth: the systematics of helium–neon isotopes inmantle-derived samples. Lithos 30, 257–265.

Honda, S., Yuen, D.A., Balachandar, S., Reuteler, D., 1993c.Three-dimensional instabilities of mantle convection with mul-tiple phase transitions. Science 259, 1308–1311.

Honnorez, J., Bonatti, E., Emilianii, C., Bronnimann, P., Furrer,M.A., Meyerhoff, A.A., 1975. Mesozoic limestone from theVema offset zone. Earth and Planetary Science Letters 26,8–12.

Hooft, E.E., Detrick, R.S., 1993. The role of density in theaccumulation of basaltic melts at mid-ocean ridges. Geophysi-cal Research Letters 20, 423–426.

Houseman, G.A., McKenzie, D.P., Molnar, P., 1981. Convectiveinstability of a thickened boundary layer and its relevance forthe thermal evolution of continental convergence belts. Journalof Geophysical Research 86, 6115–6132.

Huang, P.Y., Solomon, S.C., 1988. Centroid depths of mid-oceanridge earthquakes. Journal of Geophysical Research 93,13455–13477.

Hubbert, M.K., 1937. Theory of scale models as applied to thestudy of geologic structures. Geological Society of America

w xBulletin 48, 1459–1520 see p. 1499 footnote .Humler, E., Thirot, J.L., Montagner, J.P., 1993. Global correla-

tions of mid-ocean ridge basalt chemistry with seismic tomo-graphic images. Nature 364, 225–228.

Humler, E., Langmuir, C., Daux, V., 1999. Deth vs. age: newperspectives from the chemical compositions of ancient crust.Earth and Planetary Science Letters 173, 7–23.

Humphreys, E.D., Clayton, R.W., Hager, B.H., 1984. A tomo-graphic image of structure beneath southern California. Geo-physical Research Letters 11, 625–627.

Humphreys, E.D., Dueker, K.G., Rasmussen, J.R., 1989. Thedynamic role of the mantle on western U.S. tectonic activity.Geological Sociaty of America, Abstracts with Programs 21Ž .6 , 80–81.

Hung, S.-H., Forsyth, D.W., 1997. Upper mantle and crustalanisotropy in young oceanic lithosphere constrained from

w x Žshear-wave splitting abs. . Eos Transactions, American Geo-. Ž .physical Union 78 46 , 688.

Hunter, R.H., McKenzie, D.P., 1989. The equilibrium geometry ofcarbonate melts in rocks of mantle composition. Earth andPlanetary Science Letters 92, 347–356.

Hurley, P.M., 1966. K–Ar dating of sediments. In: Schaeffer,Ž .O.A., Zahringer, J. Eds. , Potassium–Argon Dating.

Springer-Verlag, New York, pp. 134–151.Inoue, H., Fukao, Y., Tanabe, K., Ogata, Y., 1990. Whole mantle

P-wave travel time tomography. Physics of the Earth andPlanetary Interiors 59, 294–328.

Ito, G., Clift, P.D., 1998. Subsidence and growth of PacificCretaceous plateaus. Earth and Planetary Science Letters 161,85–100.

Iwamori, H., 1998. Transportation of H O and melting in subduc-2

tion zones. Earth and Planetary Science Letters 160, 65–80.Jambon, A., Weber, H.W., Begemann, F., 1985. Helium and

argon from an Atlantic MORB glass: concentration, distribu-tion and isotopic composition. Earth and Planetary ScienceLetters 73, 255–267.

Janney, P.E., Castillo, P.R., 1999. Isotope geochemistry of theDarwin Rise seamounts and the nature of long-term dynamicsbeneath the south central Pacific. Journal of GeophysicalResearch 104, 10571–10589.

Janney, P.E., Castillo, P.R., 2000. Trace element and isotopicevidence from early Atlantic Ocean crust for deep mantle

Ž .influence. EOS 81 19 , S419.Jaroslow, G.E., Tucholke, B.E., 1994. Mesozoic–Cenozoic sedi-

mentation in the Kane fracture zone, western North Atlantic.Geological Society of America Bulletin 106, 319–337.

Javoy, M., Pineau, F., Iijima, I., 1978. Experimental determinationof the isotope fractionation between gaseous CO and carbon2

dissolved in tholeitic magma: a preliminary study. Contribu-tions to Mineralogy and Petrology 67, 35–39.

Jessop, A.M., 1990. Thermal Geophysics. Elsevier, Amsterdam.Ji, S., Zhao, X., Francis, D., 1994. Calibration of shear-wave

splitting in subcontinental upper mantle beneath active oro-genic belts using xenoliths from Canadian Cordillera andAlaska. Tectonophysics 239, 1–27.

Jones, B.F., Eugster, H.P., Rettig, S.L., 1977. Hydrochemistry ofthe Lake Magadi basin, Kenya. Geochimica et CosmochimicaActa 41, 53–72.

Jonsson, S., Einarsson, P., Sigmundsson, F., 1995. Extensionacross a divergent plate boundary: the Eastern rift zone inSouth Iceland, 1967–1994, observed with GPS and electronicdistance measurements. Journal of Geophysical Research 102,11913–11929.

Jordan, T.H., 1975. The continental tectosphere. Reviews of Geo-Ž .physics and Space Physics 13 3 , 1–12, Special Issue.

Jung, H., Karato, S., 1997. Water, partial melting and the origin ofw xseismic low velocity and high attenuation zone abs. . Eos

Ž . Ž .Transactions American Geophysical Union 78 46 , 731,Supplement.

Julian, B.R., Miller, A., Foulger, G.R., 1998. Non-double coupleearthquakes. 1. Theory. Reviews of Geophysics 36, 525–549.

Juteau, T., Cannat, M., Lagabrielle, Y., 1990. Serpentinized peri-dotites in the upper oceanic crust away from transform zones.Ocean Drilling Program, Scientific Results, Legs 106r109:College Station, pp. 303–308.

Kamenetsky, V.S., 1996. Methodology for the study of meltinclusions in Cr-spinel and implications for parental melts ofMORB. Earth and Planetary Science Letters 142, 479–486.

Ž .Kampunzu, A.B., Lubala, R.T. Eds. , 1991. Magmatism in Ex-tensional Structural Settings: The Phanerozoic African Plate.Springer-Verlag.

Kampunzu, A.B., Popoff, M., 1991. Distribution of the mainProterozoic African rifts and associated magmatism: Introduc-

Ž .tory notes. In: Kampunzu, A.B., Lubala, R.T. Eds. , Magma-tism in extensional structural settings, the Phanerozoic AfricanPlate. Springer Verlag, Berlin, pp. 2–10.

Kanasewich, R., Ellis, R.M., Chapman, C.H., Gutowski, P.R.,1973. Seismic array evidence of a core-boundary source forthe Hawaiian linear volcanic chain. Journal of GeophysicalResearch 78, 1361–1371.

Kaneoka, I., Takeoka, N., Upton, B.G.J., 1986. Noble gas system-


atics in basalts and a dunite nodule from Reunion and GrandComoro Islands, Indian Ocean. Chemical Geology 59, 35–42.

Karato, S.-I., 1987. Seismic anisotropy due to lattice-preferredŽ .orientation in minerals. In: Manghnani, M.H., Syono, Y. Eds. ,

High Pressure Research in Mineral Physics. American Geo-physical Union, Geophysical Monograph, vol. 39, pp. 455–471.

Karato, S., 1990. The role of hydrogen on the electrical conductiv-ity of the upper mantle. Nature 347, 272–273.

Karato, S.-I., Jung, H., 1998. Water, partial melting and the originof the seismic low velocity and high attenuation zone in theupper mantle. Earth and Planetary Science Letters 157, 193–207.

Ž .Karato, S.-I., Toriumi, M. Eds. , Rheology of Solids and of theEarth. Oxford Univ. Press, Oxford, UK.

Katzman, R., Zhao, L., Jordan, T.H., 1998a. High-resolution,two-dimensional vertical tomography of the central Pacificmantle using ScS reverberations and frequency-dependenttravel times. Journal of Geophysical Research 103, 17933–17971.

Katzman, R., Jordan, T.H., Zhao, L., 1998b. The strange signaturew x Žof hotspot swells in the Central Pacific abs . Eos Transac-

. Ž .tions American Geophysical Union 79 45 , 240, Fall Supple-ment.

Kawasaki, I., 1986. Azimuthaly anisotropic model of the oceanicupper mantle. Physics of Earth and Planetary Interiors 43,1–21.

Kearey, P., Vine, F.J., 1990. Global Tectonics. Blackwell.Ž .Keating, B.H., Fryer, P., Batista, R., Boehlert, W. Eds. , 1987.

Seamounts, Islands and Atolls. American Geophysical Union,wGeophysical Monograph, vol. 43, 405 pp. see papers by Okal

xand Batista, Fornari et al., Allan et al. .Keith, M.L., 1972. Ocean floor convergence: a contrary view of

global tectonics. Journal of Geology 80, 249–276.Keith, M.L., 1982. Violent volcanism, stagnant oceans, and some

inferences regarding petroleum, strata-bound ores and massextinctions. Geochimica et Cosmochimica Acta 46, 2621–2637.

Keith, M.L., 1993. Geodynamics and mantle flow: an alternativeŽEarth model. Earth Science Reviews 33, 153–337 also in

.hard cover Elsevier, Amsterdam .Keith, M.L., 1994. Dynamic surface topography: a new interpreta-

tion based on mantle flow models derived from seismic to-mography: comment. Geophysical Research Letters 21, 845–846.

Keith, C.M., Crampin, S., 1977. Seismic body waves in anisotropicmedia: synthetic seismograms. Geophysical Journal Royal As-tronomical Society 49, 225–243.

Keller, W.R., Anderson, D.L., Clayton, R.W., 1997. Resolution ofw xseismic structure in the mantle beneath Iceland abs. . Eos

Ž . Ž .Transactions American Geophysical Union 78 46 , 500.Kelley, D.S., Hoering, T.C., Frantz, J.D., 1994. Methan-rich fluids

in gabbroic rocks from the southwest Indian Ridge: Resultsfrom mass spectrometric analyses. EOS 75, 657.

Kellogg, J.B., O’Connell, R.J., 1998. Toroidal motion and mixingw x Žin three dimensions abs. . Eos Transactions, American Geo-

. Ž .physical Union 79 17 , 334, Supplement.Kemp, A.E., Pearce, R.B., Koizumi, I., Pike, J., Rance, S.J., 1999.

The role of mat-forming diatoms in formation of the Mediter-Ž .ranean sapropels. Nature London 398, 57–58.

Kendall, J.M., 1994. Teleseismic arrivals at a mid-ocean ridge:effects of mantle melt and anisotropy. Geophysical ResearchLetters 21, 301–304.

Kent, D.V., Gradstein, F.M., 1986. Jurassic to Recent chronology.Ž .In: Vogt, P.R., Tucholke, B.E. Eds. , The Western North

Atlantic Region. Geological Society of America, Geology ofNorth America, vol. M, pp. 45–50, Plate 1.

Kern, H., 1993. P- and S-wave anisotropy and shear-wave split-ting at P and T of possible mantle rocks and their relation torock fabric. Physics of the Earth and Planetary Interiors 78,245–256.

Kikawa, E., Ozawa, K., 1992. Contribution of oceanic gabbros tosea-floor spreading magnetic anomalies. Science 258, 769–799.

King, A.J., Waggoner, D.G., Garcia, M.O., 1993. Geochemistryand petrology of basalts from Leg 136, Central Pacific Ocean.Proceedings of the Ocean Drilling Program, Scientific Results136, 107–118.

Kingsley, R.H., Schilling, J.G., 1995. Carbon in Mid-AtlanticRidge basalt glasses from 288N to 638N: evidence for acarbon-enriched Azores mantle plume. Earth and PlanetaryScience Letters 129, 31–53.

Klein, F.W., Einarsson, P., Wyss, M., 1977. The ReykjanesPeninsula, Iceland, earthquake swarm of September, 1972 andits tectonic significance. Journal of Geophysical Research 82,865–888.

Klinkhammer, G.P., Palmer, M.R., 1991. Uranium in the oceans:where it goes and why. Geochimica et Cosmochimica Acta 55,1799–1806.

Klitgord, K.D., Schouten, H., 1986. Plate kinematics of the centralŽ .Atlantic. In: Vogt, P.R., Tucholke, B.E. Eds. , The Western

North Atlantic Region. Geological Society of America, Geol-ogy of North America, vol. M, pp. 351–378.

Knox, R.P., Nyblade, A.A., Langston, C.A., 1998. Upper mantleS-velocities beneath Afar and western Saudi Arabia fromRayleigh wave dispersion. Geophysical Research Letters 25,4233–4236.

Koch, D.M., Ribe, N.M., 1989. The effect of lateral viscosityvariations on surface observables. Geophysical Research Let-ters 16, 535–538.

Koga, K.T., Shimizu, M., Grove, T.L., 1998. Experimental deter-mination of spinel-garnet lherzolite stability in the mantlew x Ž .abs. . Eos Transactions American Geophysical Union 79Ž .45 , 1005, Fall Supplement.

Kogiso, T., Hirose, K., Takahashi, E., 1998. Melting experimentson homogeneous mixtures of peridotite and basalt: applicationto the genesis of ocean island basalt. Earth and PlanetaryScience Letters 162, 45–61.

Kohlstedt, D.L., Evans, B., Mackwell, S.J., 1995. Strength of thelithosphere: constraints imposed by laboratory experiments.Journal of Geophysical Research 100, 17587–17602.

Kong, L.S., Detrick, R.S., Fox, P.J., Mayer, L.A., Ryan, W.B.,1988. The morphology and tectonics of the MARC area from

ŽSea Beam and Sea MARC observations Mid-Atlantic Ridge,.238N . Marine Geophysical Research 10, 59–90.

Kumazawa, M., Anderson, O.L., 1969. Elastic moduli, pressure


derivatives and temperature derivatives of single crystalolivine. Journal of Geophysical Research 74, 5961–5968.

Kuo, B.Y., Forsyth, D.W., Wysession, M., 1987. Lateral hetero-geneity and azimuthal anisotropy in the N. Atlantic from SS-Sdifferential travel times. Journal of Geophysical Research 92,6421–6436.

Kurz, M.D., 1986a. Cosmogenic helium in a terrestrial igneousrock. Nature 320, 435–439.

Kurz, M.D., 1986b. In situ production of terrestrial cosmogenichelium and some applications to geochronology. Geochimicaet Cosmochimica Acta 50, 2855–2862.

Kurz, M.D., 1993. Mantle heterogeneity beneath oceanic islands:some inferences from isotopes. Proceedings of the RoyalSociety, London A 342, 91–103.

Kurz, M.D., Kammer, D.P., 1991. Isotopic evolution of MaunaLoa. Earth and Planetary Science Letters 103, 257–269.

Kurz, M.D., Jenkins, W.J., Hart, S.R., 1982a. Helium isotopesystematics of oceanic islands: implications for mantle hetero-geneity. Nature 297, 43–47.

Kurz, M.D., Jenkins, W.J., Schilling, J.G., Hart, S.R., 1982b.Helium isotope variations in the mantle beneath the centralNorth Atlantic. Earth and Planetary Science Letters 58, 1–14.

Kurz, M.D., Jenkins, W.J., Hart, S.R., Clague, D., 1983. Heliumisotopic variations in Loihi Seamount and the Island of Hawaii.Earth and Planetary Science Letters 66, 388–406.

Kurz, M.D., Garcia, M.O., Frey, F.A., O’Brien, P.A., 1987.Temporal helium isotopic variations within Hawaiian volca-noes: Basalts from Mauna Loa and Haleakala. Geochimica etCosmochimica Acta 51, 2905–2914.

Kurz, M.D., Colodner, D., Trull, T.W., Moore, R., O’Brien, K.,1990. Cosmic ray exposure dating with in-situ produced he-lium-3: results from young Hawaiian lava flows. Earth andPlanetary Science Letters 97, 177–189.

Kurz, M.D., Kenna, T.C., Kammer, D.P., Rhodes, J.M., Garcia,M.O., 1995. Isotopic evolution of Mauna Loa volcano: a viewfrom the submarine southwest rift zone. American Geophysi-cal Union, Geophysical Monograph 92, pp. 289–306.

Kurz, M.D., Kenna, T.C., Lassiter, J.C., DePaolo, D.J., 1996.Helium isotopic evolution of Mauna Kea volcano: first resultsfrom the 1-km drill core. Journal of Geophysical Research101, 11781–11791.

Kushiro, I., 1994. Melting behavior of mantle peridotites andcompositional variations of partial melts at high pressures.Mineralogical Magazine 58A, 503–504.

Kyser, T.K., Rison, W., 1992. Systematics of rare gas isotopes inbasic lavas and ultramafic xenoliths. Journal of Geophysical

Ž .Research B 87, 5611–5630.Langmuir, C.H., Klein, E.M., Plank, T., 1992. Petrologic system-

atics of mid-ocean ridge basalts: constraints on melt genera-tion beneath mid-ocean ridges. In: Morgan, J.P., Blackman,

Ž .D.K., Sinton, J.M. Eds. , Mantle flow and Melt Generation atMid-ocean Ridges. American Geophysical Union, GeophysicalMonograph, vol. 71, pp. 183–280.

Langston, C.A., Snell, C., O’Gorman, C., Nyblade, A.A., Owens,T.J., 1998. Deep crustal or upper mantle earthquakes associ-

w x Žated with rifting in Tanzania abs . Eos Transactions Ameri-. Ž .can Geophysical Union 79 17 , 215, Supplement.

Ž .Larson, R.L., 1990. The Pacific Jurassic, at last. Maritimes 34 3 ,1–3.

Larson, R.L., 1991. Latest pulse of Earth: evidence for a mid-Cretaceous superplume. Geology 19, 547–550.

Larson, R.L., ODP Leg 144 Shipboard Scientific Party, 1993.Highly permeable and layered Jurassic oceanic crust in theWestern Pacific. Earth and Planetary Science Letters 119,71–83.

Lassiter, J.C., Hauri, E.H., 1998. Osmium isotope variations inHawaiian lavas: evidence for recycled oceanic lithosphere inthe Hawaiian plume. Earth and Planetary Science Letters 164,483–496.

Lassiter, J.C., DePaulo, D.J., Tatsumoto, M., 1996. Isotopic evolu-tion of Mauna Kea volcano: initial results from the Hawaiiscientific drilling project. Journal of Geophysical Research101, 11769–11780.

Laubscher, H.P., 1974. The tectonics of subduction in the AlpineŽsystem. Memoria della Societa Geologica Italiana XXI Suppl.

.2 , 275–283.Lee, D.-K., Grand, S.P., 1996. Depth of the upper mantle disconti-

nuities beneath the East Pacific Rise. Geophysical ResearchLetters 23, 3369–3372.

LeFort, P., 1975. Himalayas: the collided range. American Journalof Science 275A, 1–44.

Leg 152 Shipboard Party, 1994. Drilling unearths Afire and iceB atw x Žthe S.E. Greenland Margin abs. . Eos Transactions, Ameri-

. Ž .can Geophysical Union 75 35 , 401–406.LeHuray, A.P., Church, S.E., Koski, R.A., Bouse, R.M., 1988. Pb

isotopes in sulfides from mid-ocean ridge hydrothermal sites.Geology 16, 362–365.

Lentz, D.R., 1998. A re-examination of carbonatite genesis; thew xrole of volatile fluxing in limestonermarble melting abs .

Geological Society of America, Abstracts with Programs 30Ž .7 , 24.

Leung, I.S., Li, Y.L., Han, Z.G., 1994. Metasomatized olivine,w x Žgarnet and diopside entrapped in diamonds abs. . Eos Trans-

. Ž .actions, American Geophysical Union 75 16 , 192, Supple-ment.

Lewis, C., Smith, A.D., 1997. Remixing of subducted crust intothe depleted mantle MORB, not plume source: a vindication ofthe marble cake model from Nd–Hf isotopes in komatiites and

w x ŽMORB abs. . Eos Transactions, American Geophysical. Ž .Union 78 46 , 838, Supplement.

Ligi, M., Bonatti, E., Bratti, A., Carrara, G., Gasperini, L., Vera,E., 1994. Seismic reflection images of uplifted lithosphere at

w x Žthe Vema fracture zone abs. . Eos Transactions, American. Ž .Geophysical Union 75 16 , 363, Supplement.

Lister, C.R.B., 1972. On the thermal balance of a mid-oceanridge. Geophysical Journal of the Royal Astronomical Society26, 515–535.

Lithgow-Bertelloni, C., Silver, P.G., 1998. Dynamic topography,plate driving forces and the African superswell. Nature 395,269–272.

Liu, C.-S., Curray, J.R., McDonald, J.M., 1983. New constraintson the tectonic evolution of the eastern Indian Ocean. Earthand Planetary Science Letters 65, 331–342.

Liu, J., Bohlen, S.R., Ernst, W.G., 1996. Stability of hydrous


phases in subducting oceanic crust. Earth and Planetary Sci-ence Letters 143, 161–171.

Macdonald, K.C., 1982. Mid-ocean ridges: fine scale tectonic,volcanic and hydrothermal processes within the plate bound-ary zone. Annual Review of Earth and Planetary Science 10,155–190.

Macdonald, K.C., 1986. The crest of the Mid-Atlantic Ridge:models for crustal generation processes and tectonics. In:

Ž .Vogt, P.R., Tucholke, B.E. Eds. , The Western North AtlanticRegion. Geology of North America, vol. M, Geological Soci-ety of America, Boulder, CO, pp. 51–68.

Macdonald, K.C., Luyendyk, B.P., 1977. Deep-tow studies of thestructure of the Mid-Atlantic Ridge near 378N. GeologicalSociety of America Bulletin 88, 621–636.

Macdonald, K.C., Sempere, J.C., Fox, P.J., 1986a. The debateconcerning overlapping spreading centers and mid-ocean ridgeprocesses. Journal of Geophysical Research 91, 10501–10511Ž .cf. Lonsdale, op. cit., p. 10,493 .

Macdonald, K.C., Castillo, D.A., Miller, S.P., Fox, P.J., Kastens,K.A., Bonatti, E., 1986b. Deep-tow studies of the Vemafracture zone: 1. Tectonics of a slow-slipping transform faultand its intersection with the Mid-Atlantic Ridge. Journal ofGeophysical Research 91, 3334–3354.

Macdonald, K.C. et al., 1988. A new view of the mid-ocean ridgefrom the behavior of ridge-axis discontinuities. Nature 335,217–225.

Machetel, P., Weber, P., 1991. Intermittent layered convection ina model mantle with an endothermic phase transition at 670km. Nature 350, 55–57.

Magde, L.S., Dick, H.J.B., Hart, S.R., 1995. Tectonics, alterationand fractal distribution of hydrothermal veins in the lowerocean crust. Earth and Planetary Science Letters 129, 103–119.

Maia, M., Diament, M., 1991. Analysis of the altimetric geoid invarious wavebands in the central Pacific Ocean. Tectono-physics 190, 133–153.

Mainprice, D., Silver, P.G., 1993. Interpretation of SKS wavesusing samples from subcontinental lithosphere. Physics ofEarth and Planetary Interiors 78, 257–280.

Malinconico, L.L., 1989. Crustal thickness estimates for the west-ern Himalaya. Geological Society of America, Special Paper232, 237–242.

Malinvemo, A., 1988. Non steady-state topography of the Mid-w x ŽAtlantic Ridge abs. . Eos Transactions, American Geophysi-

. Ž .cal Union 69 44 , 1422.Margaritz, M., Turner, P., 1982. Carbon cycle changes of the

Zechstein sea, isotopic transition zone in the Marl Slate.Nature 297, 389–390.

Martin, C.E., Carlson, R.W., Shirey, S.B., Frey, F.A., Chen,C.-H., 1994. Os isotopic variation in basalts from Haleakalavolcano, Maui, Hawaii. Earth and Planetary Science Letters128, 287–301.

Marty, B., Humbert, F., 1997. Nitrogen and argon isotopes inoceanic basalts. Earth and Planetary Science Letters 152,101–112.

Marty, B., Ozima, M., 1986. Noble gas distribution in oceanicbasalt glasses. Geochimica et Cosmochimica Acta 50, 1093–1097.

Marty, B., Tolstikhin, I.N., 1998. CO fluxes from mid-ocean2

ridges, arcs and plumes. Chemical Geology 145, 233–248.Marty, B., Pik, R., Gezahegri, Y., 1996. Helium isotopic varia-

tions in Ethiopian plume lavas, nature of magmatic sourcesand limit on lower mantle contribution. Earth and PlanetaryScience Letters 144, 223–237.

Matthews, S.W., 1973. This changing Earth. National GeographicŽ .Magazine 143 1 , 1–15.

Maurin, J.C., Guiraud, R., 1993. Basement control in the develop-ment of the early Cretaceous West and Central African riftsystem. Tectonophysics 222, 81–95.

McAllister, E., Cann, J.R., 1996. Initiation and evolution ofboundary wall faults along the Mid-Atlantic Ridge. Tectonic,Magmatic and Hydrothermal Segmentation of Mid-Ocean

Ž .Ridges, Geological Society Special Publication London , vol.118, pp. 29–48.

McCourt, S., Vearncombe, J.R., 1987. Shear zones bounding thecentral zone of the Limpopo mobile belt, southern Africa.Journal of Structural Geology 9, 127–137.

McDonough, D.T., Fountain, D.M., 1993. P-wave anisotropy ofmylonite and infrastructural rocks from a Cordilleran corecomplex. Physics of Earth and Planetary Interiors 78, 319–336.

McDougall, I., Honda, M., 1998. Primordial solar noble gasŽ .component in the Earth. In: Jackson, I. Ed. , The Earth’s

Mantle: Composition, Structure and Evolution, pp. 159–187.McDougall, I., Kristjansson, L., Saemondsson, K., 1984. Magne-

tostratigraphy and chronology of northwest Iceland. Journal ofGeophysical Research 89, 7029–7060.

McKenzie, D., 1979. Finite deformation during fluid flow. Geo-physical Journal, The Royal Astronomical Society 58, 689–715.

McKenzie, D., Roberts, J., Weiss, M., 1973. Numerical models ofconvection in the Earth’s mantle. Tectonophysics 19, 89–103.

McNamara, D.E., Owens, T.J., Silver, P.G., Wu, F.T., 1994.Shear wave anisotropy beneath the Tibetan Plateau. Journal ofGeophysical Research 99, 13655–13665.

Mei, S., Kohlstedt, D.L., 1997. Effect of water on plastic deforma-w x Žtion of olivine aggregates abs. . Eos Transactions, American

. Ž .Geophysical Union 78 46 , 730, Supplement.Melson, W.G., O’Hearn, T., 1986. Zero age variations in the

composition of abyssal volcanic rocks along the axial zone ofŽ .the Mid-Atlantic Ridge. In: Vogt, P.R., Tucholke, B.E. Eds. ,

The Western North Atlantic Region. Geological Society ofAmerica, Geology of North America, vol. M, pp. 117–136.

Melson, W.G., Hart, S.R., Thompson, G., 1972. St. Paul’s Rocks,equatorial Atlantic: petrogenesis, radiometric ages and impli-cations on sea-floor spreading. Geological Society of America,Memoir 132, 241–272.

MELT seismic team, 1998. Imaging the deep seismic structurebeneath a mid-ocean ridge: the MELT experiment. Science

Ž280, 1215–1218 see related papers, 1219–1238, by Toomey.et al., Webb and Forsyth, Wolfe and Solomon .

Menard, H.W., 1964. Marine Geology of the Pacific. McGraw-Hill, New York, 271 pp.

Menard, H.W., Chase, T.E., 1970. Fracture zones. In: Maxwell,Ž .A.E. Ed. , 1970. The Sea, vol. 4, pp. 421–443.

Menke, W., Bransdottir, B., Einarsson, P., 1995. The role of


w xKrafla volcano in crustal formation in Iceland abs. . EosŽ . Ž .Transactions, American Geophysical Union 76 46 , 582,Supplement.

Michael, P.J., Schilling, J.G., 1989. Chlorine in mid-ocean ridgemagmas: evidence for assimilation of seawater-influencedcomponents. Geochimica et Cosmochimica Acta 53, 3131–3143.

Miller, D.M., Busby, C., 1995. Jurassic Magmatism and Tectonicsof the North American Cordillera. Geological Society ofAmerica, Boulder, CO, Special Paper 299, 432 pp.

Miller, A.D., Julian, B.R., Foulger, G.R., 1998. Three dimensionalseismic structure and moment tensors of non-double coupleearthquakes at the Hengill Grensdalur volcanic complex, Ice-land. Geophysical Journal International 133, 309–325.

Milton, C., 1968. The Anatro-carbonatiteB lava of Oldoinyo Lengai,w xTanzania abs . Geological Society of America, Special Paper

121, 202.Milton, C., 1989. Oldoinyo Lengai natrocarbonatite lava: its his-

w xtory abs . 28th International Geological Congress, Washing-ton, DC, vol. 2, p. 441.

Mitchell, N., Escartin, J., Allerton, S., 1998. Detachment faults atw x Žmid-ocean ridges abs. . Eos Transactions, American Geo-

. Ž .physical Union 79 10 , 127.Molnar, P., 1988. A review of geophysical constraints on the deep

structure of the Tibetan plateau, the Himalaya and the Karako-rum. Philosophical Transactions of the Royal Society of Lon-don, A 326, 33–88.

Montagner, J.P., Anderson, D.L., 1989. Constrained referencemantle model. Physics of the Earth and Planetary Interiors 58,205–227.

Montagner, J.P., Nataf, H.C., 1986. A simple method for invertingthe azimuthal anisotropy of surface waves. Journal of Geo-physical Research 91, 511–520.

Montagner, J.P., Nataf, H.C., 1988. Vectorial tomography: I.Theory. Geophysical Journal of the Royal Astronomical Soci-ety 94, 295–307.

Montagner, J.P., Tanimoto, T., 1990. Global anisotropy in theupper mantle inferred from the regionalization of phase veloci-ties. Journal of Geophysical Research 95, 4797–4819.

Moore, J.G., Batchelder, J.N., Cunningham, C.G., 1977. CO -filled2

vesicles in mid-ocean ridge basalt. Journal of Volcanology andGeothermal Research 2, 309–327.

Moos, D., Zoback, M.D., 1990. Utilization of well bore failure toconstrain the orientation of crustal stresses. Journal of Geo-physical Research 95, 9305–9325.

Moreira, M., Kunz, J., Allegre, C.J., 1998. Rare gas systematics inpopping rock: isotopic and elemental compositions in theupper mantle. Science 279, 1178–1181.

Moreira, M., Staudacher, T., Sarda, P., Schilling, J.-G., Allegre,C.J., 1991. A primitive plume neon component in MORB: TheShona ridge anomaly, South Atlantic. Earth and PlanetaryScience Letters 133, 367–377.

Ž .Morgan, J.P., Blackman, D.K., Sinton, J.M. Eds. , Mantle Flowand Melt Generation at Mid-ocean Ridges. American Geo-physical Union, Geophysical Monograph, vol. 71.

Morgan, W.J., 1998. Personal reflections on the origin of plate

w x Žtectonics abs. . Eos Transactions, American Geophysical. Ž .Union 79 45 , F15.

Morris, G.B., Raitt, R.W., Shor, G.C., 1969. Velocity anisotropyand delay time map of the mantle near Hawaii. Journal ofGeophysical Research 74, 4300–4316.

Morris, E., Dettrick, R.S., Minshull, T.A., White, R.S., Su, W.,Buhl, P., 1993. Seismic structure of oceanic crust in thewestern North Atlantic. Journal of Geophysical Research 98,13879–13903.

Mottl, M.J., Wheat, C.J., 1994. Hydrothermal circulation throughmid-ocean ridge flanks: fluxes of heat and magnesium.Geochimica et Cosmochimica Acta 58, 2225–2237.

Mueller, S., 1989. Deep-reaching dynamic processes in the Alps.Ž .In: Coward, M.P. Ed. , Alpine Tectonics. Geological SocietyŽ .Special Publication London , vol. 45, pp. 303–328.

Mutter, J.G., Karson, J.A. et al., 1992. Structural processes atslow-spreading ridges. Science 257, 627–634.

Mutter, J.C., Talwani, M., Stoffa, P.L., 1984. Evidence for a thickoceanic crust adjacent to the Norwegian margin. Journal ofGeophysical Research 89, 483–502.

Myers, J.D., Johnson, A.D., 1996. Phase equilibria constraints onmodels of subduction zone magmatism. In: Bebout, G.E.,

Ž .Scholl, D.W., Kirby, S.H., Platt, J.P. Eds. , Subduction Topto Bottom. American Geophysical Union, Geophysical Mono-graph, vol. 96, pp. 229–249.

Ž .Mysen, B.O. Ed. , 1987. Magmatic Processes: PhysicochemicalPrinciples: A Volume in Honor of Hatten S. Yoder. Geochem-ical Society Special Publication, vol. 1.

Navrotsky, A., 1993. How much do we know about mantlethermochemistry? Science 261, 168–169.

Neumann, G.A., Forsyth, D.W., 1993. The paradox of the axialprofile: isostatic compensation along the axis of the Mid-Atlantic Ridge. Journal of Geophysical Research 98, 17891–17910.

Nicolas, A., 1993. Why fast polarization directions of SKS seis-mic waves are parallel to mountain belts. Physics of Earth andPlanetary Interiors 78, 337–342.

Nicolas, A., Christensen, N.I., 1987. Formation of anisotropy inupper mantle peridotites: a review. In: Fuchs, K., Froidevaux,

Ž .C. Eds. , Composition, Structure and Dynamics of the Litho-sphere–Asthenosphere System. American Geophysical Union,Geodynamics Series, vol. 16, pp. 111–123.

Nicolas, A., Boudier, F., Boulier, A.M., 1973. Mechanisms offlow in naturally and experimentally deformed peridotites.American Journal of Science 273, 853–870.

Niedermann, S., Bach, W., Erzinger, J., 1997. Noble gas evidencefor a lower mantle component in MORB from the southernEast Pacific Rise: decoupling of helium and neon isotopesystematics. Geochimica et Cosmochimica Acta 61, 2697–2713.

Nielsen, R.L., Walker, A., Christie, D.M., Sours-Page, R., 1995.Local, regional and global variations in incompatible Aoveren-

Ž .richmentB in MORB lavas. EOS 76 46 , 795.Nishimura, C.E., Forsyth, D.W., 1989. The anisotropic structure

of the upper mantle in the Pacific. Geophysical Journal Inter-national 96, 203–229.


Ž .Nixon, P.H. Ed. , Mantle Xenoliths. Wiley-Interscience, NewYork.

Nyblade, A.A., Pollack, H.N., 1993. A global analysis of heatflow from Precambrian terrains. Journal of Geophysical Re-search 98, 12207–12218.

Nyblade, A.A., Pollack, H.N., Jones, D.L., Podmore, F.,Mushayandebvu, M., 1990. Terrestrial heat flow in East andsouthern Africa. Journal of Geophysical Research 95, 17371–17384.

O’Connor, J.M., Stoffers, P., McWilliams, M.O., 1995. Time-space mapping of Easter chain volcanism. Earth and PlanetaryScience Letters 136, 197–212.

Okal, E.A., Batista, R., 1987. Hotspots: The first 25 years. In:Ž .Keating, B.H. Ed. , Seamounts, Islands and Atolls. American

Geophysical Union, Geophysical Monograph, vol. 43, pp.1–11.

Olafsson, M., Eggler, D.H. et al., 1983. Phase relations of amphi-bole, phlogopite and carbonate peridotite. Earth and PlanetaryScience Letters 64, 305–315.

O’Neil, J.R., Hay, R.L., 1973. Oxygen isotope ratios in chertsassociated with saline lake deposits of East Africa. Earth andPlanetary Science Letters 19, 257–266.

Ozima, M., Saito, K., Matsuda, J., Zashu, S., Aramaki, S., Shido,F., 1976. Additional evidence of ancient rocks on the Mid-Atlantic Ridge. Tectonophysics 31, 59–71.

Palacz, Z.A., Saunders, A.D., 1986. Coupled trace element andisotope enrichment in Cook–Austral–Samoa Islands. Earthand Planetary Science Letters 79, 270–280.

Palmason, G., 1986. Model of crustal formation in Iceland andapplication to submarine mid-ocean ridges. In: Vogt, P.E.,

Ž .Tucholke, B.E. Eds. , The Western North Atlantic Region.Geological Society of America, Geology of North America,vol. M, pp. 87–97.

Palmason, G. et al., 1979. The Iceland crust: evidence fromdrillhole data. In: Talwani, M., Harrison, C.G., Hayes, D.E.Ž .Eds. , Deep drilling results in the Atlantic Ocean crust.American Geophysical Union, Maurice Ewing Series, vol. 2,pp. 43–65.

Papale, P., 1995. Modeling the solubility of a two-componentH O-CO fluid in silicate liquids. EOS 76, 706.2 2

Pariso, J.E., Johnson, H.P., 1993. Do lower crustal rocks recordreversals of Earth’s magnetic field? Journal of GeophysicalResearch 98, 16013–16032.

Pariso, J.E., Sempere, J.C., Rommevaux, C., 1995. Temporal andspatial variations in crustal accretion along the Mid-AtlanticRidge. Journal of Geophysical Research 100, 17781–17794.

Parsons, B., McKenzie, D., 1978. Mantle convection and thethermal structure of the plates. Journal of Geophysical Re-search 83, 4485–4496.

Patterson, D.B., Honda, M., McDougall, I., 1990. Atmosphericcontamination: a possible source for heavy noble gases inbasalts from Loihi Seamount, Hawaii. Geophysical ResearchLetters 17, 705–708.

Patterson, D.B., Honda, M., McDougall, I., 1994. Deconvolutionof multiple components of neon and helium in mantle-derived

Ž .samples. In: Matsuda, J. Ed. , Noble gas geochemistry andcosmochemistry. Terra Scientific Publishing, pp. 179–189.

Pearson, D.G., Shirey, S.B., Carlson, R.W., Boyd, F.R.,Pokhilenko, N.P., Shimizu, N., 1995a. Re–Os, Sm–Nd andRb–Sr isotopic evidence for thick Archean lithospheric mantlebeneath the Siberian craton. Geochimica et CosmochimicaActa 59, 959–977.

Pearson, N.J., O’Reilly, S.Y., Griffin, W.L., 1995b. The crust-mantle boundary beneath cratons and craton margins: A tran-sect across the southwest margin of the Kaapvaal craton.Lithos 36, 257–287.

Pegram, W.J., Krishnaswami, S., Ravizza, G., Turekian, K.K.,1992. The record of seawater 187Osr186Os variation throughthe Cenozoic. Earth and Planetary Science Letters 113, 569–576.

Perfit, M., 1999. Earth’s oceanic crust. In: Marshall, C.P., Fair-Ž .bridge, W. Eds. , Encyclopedia of Geochemistry. Kluwer

Academic Publishers, pp. 179–182.Peuker-Ehrenbrink, B., Blum, J.D., Ravizza, G., 1997. The effects

of surficial processes on the marine Os and Sr isotope records.Ž .Geological Society of America Abstracts with Program 29 6 ,

169.Ž .Phipps-Morgan, J., Blackman, D.K., Sinton, J.M. Eds. , Mantle

Flow and Melt Generation at Mid-Ocean Ridges. AmericanGeophysical Union, Geophysical Monograph, vol. 71.

Pilot, J., Werner, C.D., Haubrich, F., Baumann, N., 1998. Paleo-zoic and Proterozoic zircons from the Mid-Atlantic Ridge.Nature 393, 676–677.

Pineau, F., Javoy, M., 1994. Strong degassing at ridge crests:dissolved carbon and water in basalt glasses, Mid-AtlanticRidge. Earth and Planetary Science Letters 123, 179–198.

Piper, D.Z., Isaacs, C.M., Dean, W.E., 1994. Minor elementinterpretation of deposition: modern sediment and ancientshales. Geological Society of America, Abstracts with Pro-

Ž .grams 26 7 , 29.Plank, T., Langmuir, C.H., 1998. The chemical composition of

subducting sediment and its consequences for crust and man-tle. Chemical Geology 145, 325–394.

Platt, J.P., England, P.C., 1994. Convective removal of lithospherebeneath mountain belts: thermal and mechanical conse-quences. American Journal of Science 294, 307–336.

Polet, J. et al., 1996. Shear wave anisotropy beneath the centralAndes from the BANJO, SEDA and PISCO experimentsw x Ž .abs. . Eos Transactions, American Geophysical Union 77Ž .46 , 464, Supplement.

Poli, S., 1993. The amphibolite–eclogite transformation: an exper-imental study on basalt. American Journal of Science 293,1061–1107.

Pollitz, F.F., Sacks, I.S., 1996. Viscosity structure beneath north-Ž .east Iceland. Journal of Geophysical Research B 101, 17771–

17793.Popp, T., Kern, H., 1993. Thermal dehydration reactions charac-

terized by combined measurements of electrical conductivityand elastic wave velocities. Earth and Planetary Science Let-ters 120, 43–57.

Poreda, R.J., Schilling, J.G., Craig, H., 1993. Helium isotopes inEaster microplate basalts. Earth and Planetary Science Letters119, 319–329.

Ž .Pringle, M.S., Sager, W.W., Sliter, W.V., Stein, S. Eds. , The


Mesozoic Pacific, Geology, Tectonics and Volcanism. Ameri-can Geophysical Union, Monograph, vol. 77.

Prinzhofer, A., Lewin, E., Allegre, C.J., 1989. Stochastic meltingof the marble cake mantle: evidence from the East PacificRise. Earth and Planetary Science Letters 92, 189–206.

Pyle, D.M., Dawson, J.B., Ivanovich, M., 1991. Short-lived decayseries disequilibrium in the natrocarbonatite lavas of OldoinyoLengai, Tanzania: constraints on the timing of magma genesis.Earth and Planetary Science Letters 105, 378–396.

Raitt, R.W., Shor, G.G., Morris, G.B., Kirk, H.K., 1971. Mantleanisotropy in the Pacific Ocean. Tectonphysics 12, 173–186.

Ravizza, G., 1993. Variations of the 187Osr186Os ratio of seawa-ter over the past 28 million years as inferred from metallifer-ous carbonates. Earth and Planetary Science Letters 118,335–348.

Reisberg, L., Zindler, A., Marcantonio, F., White, W., Wyman,D., Weaver, B., 1993. Os isotope systematics in ocean islandbasalts. Earth and Planetary Science Letters 120, 149–167.

Rehkamper, M., Hofmann, A.W., 1997. Recycled ocean crust andsediment in Indian Ocean MORB. Earth and Planetary ScienceLetters 147, 93–106.

Reston, T.S., Ranero, C.R., 1999. The structure of Cretaceousoceanic crust of the NW Pacific: constraints on processes atfast-spreading centers. Journal of Geophysical Research 104,629–644.

Rhodes, J.M., Hart, S.R., 1995. Episodic trace element and iso-topic variations in historical Mauna Loa lavas: implications formagma and plume dynamics. American Geophysical Union,Geophysical Monograph 92, pp. 263–276.

Ribe, N.M., 1989. Seismic anisotropy and mantle flow. Journal ofGeophysical Research 94, 4213–4223.

Ribe, N.M., Yu, Y., 1991. A theory for plastic deformation andtextural evolution of olivine polycrystals. Journal of Geophysi-cal Research 96, 8325–8335.

Ricciardi, K., Abbott, D., 1993. Increased mantle convectionduring the mid-Cretaceous. EOS 74, 633.

Richards, M.A., Hager, B.H., Sleep, N.H., 1988. Dynamicallysupported geoid highs over hot spots: observations and theory.Journal of Geophysical Research 93, 7690–7708.

Richet, P., Ingrin, J., Mysen, B.O., Courtial, P., Gillet, P., 1994.Premelting effects in minerals: an experimental study. Earthand Planetary Science Letters 121, 589–600.

Richter, F.M., 1973. Convection and the large-scale circulation ofthe mantle. Journal of Geophysical Research 78, 8735–8745.

Ring, U., 1993. Kinematic history and mechanisms of superposi-tion of the Proterozoic mobile belts of eastern central Africa.

Ž .Precambrian Research 62 3 , 207–226.Richter, F.M., Parsons, B., 1975. On the interaction of two scales

of convection in the mantle. Journal of Geophysical Research80, 2529–2541.

Rison, W.A., Craig, H., 1983. Helium isotopes and mantle volatilesin Loihi Seamount and Hawaiian Island basalts and xenoliths.Earth and Planetary Science Letters 66, 407–426.

Ritsema, J., Helmberger, D., Ni, S., Wen, L., 1998. New seismicconstraints on a large thermal upwelling in Earth’s mantle

w x Žbeneath Africa abs. . Eos Transactions, American Geophysi-. Ž .cal Union 79 17 , 213, Supplement.

Robinson, E.M., Parsons, B., Driscoll, M., 1988. Effect of a

shallow low viscosity zone on mantle flow at fracture zones.Journal of Geophysical Research 93, 25–43, 3469–3479.

Roden, M.F., Trull, T., Hart, S.R., Frey, F.A., 1994. New He, Nd,Pb and Sr isotopic constraints on the constitution of theHawaiian plume: results from Koolau volcano, Oahu, Hawaii.Geochimica et Cosmochimica Acta 58, 1431–1440.

Roeder, P.L., 1994. Chromite: from the fiery rain of chondrules tothe Kilauea lava lake. The Canadian Mineralogist 32, 729–746.

Rognon, P., Coude-Gaussen, G., Revel, M., Grousset, F.E., Pede-may, P., 1996. Holocene Saharan dust deposition in the CapeVerdes islands. Sedimentology 43, 359–366.

Rowlett, H., Forsyth, D.W., 1984. Recent faulting and mi-croearthquakes at the intersection of Vema fracture zone andthe Mid-Atlantic Ridge. Journal of Geophysical Research 89,6079–6094.

Roy-Barman, M., Allegre, C.J., 1994. 187Osr186Os of mid-oceanridge basalts and abyssal peridotites. Geochimica et Cos-mochimica Acta 58, 5043–5054.

Ž .Ruddiman, W.F. Ed. , Tectonic Uplift and Climate Change.ŽPlenum, New York, see papers by Edmond and Huh, Kump

.and Arthur .Ruppert, S., Fliedner, M.M., Zandt, G., 1998. Thin crust and

active upper mantle beneath the southern Sierra Nevada in thewestern United States. Tectonophysics 286, 237–252.

Russo, R.M., Silver, P.G., 1994. Trench-parallel flow beneath theNazca Plate from seismic anisotropy. Science 263, 1105–1111.

Rutter, E.H., Brodie, K.H., 1988. Experimental AsyntectonicBdehydration of serpentinite under conditions of controlledwater pressure. Journal of Geophysical Research 93, 4907–4932.

Ž .Ryan, M.-P. Ed. , 1994. Magmatic Systems. Academic Press,ŽSan Diego, 401 pp. see chapters by Ryan, Phipps-Morgan et

.al. .Saemundsson, K., 1986. Subaerial volcanism in the western North

Ž .Atlantic. In: Vogt, P.E., Tucholke, B.E. Eds. , The WesternNorth Atlantic Region. Geological Society of America, Geol-ogy of North America, vol. M, pp. 69–86.

Sandvol, E.A., Ni, J.F., Hearn, T.M., 1994. Seismic anisotropybeneath the Pakistan Himalayas. Geophysical Research Letters21, 1635–1639.

Sarda, P., Graham, D., 1990. Mid-ocean ridge popping rocks:implications for degassing at ridge crests. Earth and PlanetaryScience Letters 97, 268–269.

Sarda, P., Staudacher, T., Allegre, C.J., 1983. Comparative Ar-Xesystematics in the earth’s mantle. In: Pepin, R.O., O’Connell,

Ž .R.J. Eds. , Conference on planetary volatiles, Alexandria,MN. Lunar and Planetary Inst., vol. 83-01, p. 148.

Sarda, P., Staudacher, T., Allegre, C.J., 1985. 40Arr36Ar in MORBglasses: constraints on atmosphere and mantle evolution. Earthand Planetary Science Letters 72, 357–375.

Sarda, P., Staudacher, T., Allegre, C.J., 1988. Neon isotopes insubmarine basalts. Earth and Planetary Science Letters 91,73–88.

Sato, H., Sacks, I.S., 1989. Anelasticity and thermal structure ofthe oceanic upper mantle: Temperature calibration with heat

Ž .flow data. Journal of Geophysical Research B 94, 5705–5715.

Sato, H., Sacks, S.I., Murase, T., Muncill, G., Fukuyama, H.,


1989. Qp-melting temperature relation in peridotite at highpressure and temperature: Attenuation mechanism and impli-cations for mechanical properties of the upper mantle. Journalof Geophysical Research 94, 10647–10661.

Scarsi, P., Craig, H., 1996. Helium isotope ratios in Ethiopian riftbasalts. Earth and Planetary Science Letters 144, 505–516.

Schafer, K., 1979. In situ strain relief measurements in Iceland.Tectonophysics 52, 118–119.

Scherer, P., Schultz, L., Loeken, T., 1994. Weathering and atmo-Ž .spheric noble gases in chondrites. In: Matsuda, J. Ed. , Noble

Gas Geochemistry and Cosmochemistry. Terrapub, Tokyo, pp.43–53.

Schiano, P., Brick, J.-L., Allegre, C.J., 1997. Osmium–strontium–neodymium–lead isotope covariations in mid-oceanridge basalt glasses and the heterogeneity of the upper mantle.Earth and Planetary Science Letters 150, 363–379.

Schilling, J.G., 1986. Geochemical and isotopic variation alongthe Mid-Atlantic Ridge axis from 798N to 08N. In: Vogt, P.E.,

Ž .Tucholke, B.D. Eds. , The Western North Atlantic Region.Geological Society of America, Geology of North America,vol. M, pp. 137–156.

Schmidt, M.W., Poli, S., 1994. The stability of lawsonite andzoisite at high pressures: experiments to 92 kbar and implica-tions for the presence of hydrous phases in subduction zones.Earth and Planetary Science Letters 124, 105–118.

Schrauder, M., Navon, O., 1994. Carbonate and water-bearingfluids in fibrous diamonds from Botswana. Geochimica etCosmochimica Acta 58, 761–771.

Schubert, G., Sandwell, D., 1989. Crustal volume of the conti-nents and of oceanic plateaus. Earth and Planetary ScienceLetters 92, 234–246.

Sclater, J.G., Fisher, R.L., 1974. Evolution of the East CentralIndian Ocean. Bulletin of the Geological Society of America85, 683–702.

Sclater, J.G., Wixon, L., 1986. Heat flow and age in the NorthŽ .Atlantic. In: Vogt, P.R., Tucholke, B.E. Eds. , The Western

North Atlantic Region. Geological Society of America, Geol-ogy of North America, vol. M, pp. 257–270.

Searle, R., 1979. Side-scan sonar studies of North Atlantic frac-ture zones. Quarterly Journal, Geological Society of London136, 283–292.

Searle, R.C., 1983. Multiple closely spaced transform faults infast-slipping fracture zones. Geology 11, 607–610.

Sempere, J.C., Lin, J., Brown, H.S., Schouten, S., Purdy, G.M.,1993. Segmentation and morphotectonic variations along aslow-spreading center of the Mid-Atlantic Ridge. Marine Geo-

w xphysical Researches 15, 153–200 see Figs. 3, 9 and 11 .Shand, S.J., 1930. Limestone and the origin of feldspathoidal

rocks. Geological Magaine 67, 415–427.Shaw, H.R., Jackson, E.D., 1973. Linear island chains in the

Pacific: result of thermal plumes or gravitational anchors?Journal of Geophysical Research 78, 8634–8652.

Shen, Y., Forsyth, D.W., 1995. Geochemical constraints on initialand final depths of melting beneath mid-ocean ridges. Journalof Geophysical Research 100, 2211–2237.

Shen, Y., Solomon, S.C., Bjarnason, I.T., Purdy, G.M., 1996. Hotmantle transition zone beneath Iceland and the adjacent Mid-Atlantic Ridge inferred from P-to-S conversions at the 410-

and 660-km discontinuities. Geophysical Research Letters 23,3527–3530.

Shen, Y., Solomon, S.C., Bjarnason, I.T., Wolfe, C.J., 1998a.Seismic evidence for a lower mantle origin of the Iceland

Ž .plume. Nature London 395, 62–65.Shen, Y., Sheehan, A.F., Dueker, J.G., Groot-Hedlin, C., Gilbert,

H., 1998b. Mantle discontinuity structure beneath the southernEast Pacific Rise from P-to-S converted phases. Science 280,1232–1234.

Shih, X.R., Schneider, J.F., Meyer, R.P., 1991. Polarities of P andS waves and shear-wave splitting observed from the Bucara-manga nest, Colombia. Journal of Geophysical Research 96,12069–12082.

Shimamura, H., Asada, T., Suyehiro, Y., Yamada, T., Inatane, H.,1983. Longshot experiments to study velocity anisotropy inoceanic lithosphere. Physics of Earth and Planetary Interiors31, 348–362.

Shipboard Scientific Party, 1974. Site 253. Initial Reports of theDeep Sea Drilling Project, vol. 22, U.S. Government PrintingOffice, Washington, DC, pp. 153–232.

Shipboard Scientific Party, 1976. Site 384: the Cretaceous–Ter-tiary boundary, Aptian reefs and the J-anomaly ridge. In:

Ž .Tucholke, B.E., Vogt, P.R. Eds. , Initial Reports of the DeepSea Drilling Project, vol. 43, U.S. Government Printing Of-fice, Washington, DC, pp. 107–154.

Shipboard Scientific Party, 1977. Site 332. In initial reports of theDeep sea Drilling Project, vol. 37, 15–199. U.S. Govt. Print-ing Office, Washington, D.C.

Shipboard Scientific Party, et al., 1978a. Site 395. In: Melson,Ž .W.G., Rabinowitz, P.D. Eds. , Initial Reports of the Deep Sea

Drilling Project, vol. 45, p. 131.Shipboard Scientific Party, et al., 1978b. Sites 407, 408 and 409.

Ž .In: Luyendyk, B.P., Cann, J.R. Eds. , Initial Reports of theDeep-Sea Drilling Project, vol. 49, U.S. Government PrintingOffice, Washington, pp. 21–225.

Shipboard Scientific Party, 1979. Site 417. In: T. Donnelly, J.Ž .Francheteau et al. Eds. , Initial Reports of the Deep Sea

Drilling Project, 51, pp. 25, 75.Shipboard Scientific Party, et al., 1989. Site 73. In: Robinson,

Ž .P.T., Von Herzen, R. Eds. , Proceedings of the Ocean DrillingProgram, Initial Reports, vol. 118, pp. 41–57.

Shipboard Scientific Party, 1998. In: Whitmarsh, R.B., Beslier,Ž .M.O. et al. Eds. , Proceedings, Ocean Drilling Program,

Initial Reports, College Station, TX, 493 pp.Sigmundsson, F., Vadoh, H., 1996. Oblique spreading at the

Mid-Atlantic Ridge, SW Iceland, inferred from satellite radarw x Žand geologic evidence abs. . Eos Transactions, American

. Ž .Geophysical Union 77 46 , 143, Supplement.Sigmundsson, F., Vadon, H., Massonnet, D., 1997. Readjustment

of the Krafla spreading segment to crustal rifting measured bysattelit radar interfereometry. Geophysical Research Letters24, 1843–1846.

Sigurdsson, H., 1981. First order major element variation in basaltglasses from the Mid-Atlantic Ridge: 298N to 738N. Journal ofGeophysical Research 86, 9483–9502.

Sigurdsson, H., 1987. Dyke injection in Iceland. In: Halls, H.C.,Ž .Fahrig, W.F. Eds. , Mafic Dyke Swarms. Geological Associa-

tion of Canada, Special Paper, vol. 34, pp. 55–64.


Silver, P.G., Chan, W.W., 1991. Shear-wave splitting and sub-continental mantle deformation. Journal of Geophysical Re-search 96, 16429–16454.

Simiyu, S.M., Keller, G.R., 1996. An integrated analysis oflithospheric structure across the East African plateau based ongravity anomalies and recent seismic studies. Tectonophysics278, 291–313.

Smith, W., Sandwell, D., 1997. Measured and estimated seafloorŽ .topography version 4.2 . World Data Center A for Marine

Geology and Geophysics, Publication RP-1, poster 340=530.Smith, M.C., Perfit, M.R., Embley, R.W., Chadwick, W.W.,

1997. Interpretations of magmatic activity and crustal accre-tion along the Coaxial Segment and Axial Seamount NorthRift Zone, using combined acoustic and geochemical data tomap the seafloor at the scale of individual flows. EOS 78, 676.

Sobolev, A.V., Shimizu, N., 1994. The origin of typical N-MORB:the evidence from a melt inclusion study. Mineralogical Maga-zine 58A, 862–863.

Solomon, M., 1990. Subduction, arc reversal and the origin ofporphyry copper–gold deposits in island arcs. Geology 18,630–633.

Souriau, A., 1981. The upper mantle beneath Ninety-East Ridgeand Broken Ridge, Indian Ocean. Geophysical Journal of theRoyal Astronomical Society 67, 359–374.

Souriau, A., Njike-Kassala, J.D., 1993. Anisotropy beneath theNorth Pyrenean fault. Physics of Earth and Planetary Interiors78, 239–244.

Srivastava, S.P., Topscott, C.R., 1986. Plate kinematics of theŽ .North Atlantic. In: Vogt, P.M., Tucholke, B.E. Eds. , The

Western North Atlantic Region. Geology of North America,vol. M, Geological Society of America, Boulder, CO, pp.379–404.

Staudacher, T., Kurz, M.D., Allegre, C.J., 1986. New noble gasdata on glass samples from Loihi Seamount and Hualalai andon dunite samples from Loihi and Reunion island. ChemicalGeology 56, 193–205.

Staudacher, T., Sarda, P., Richardson, S.H., Allegre, C.R., Sagna,I., Dmieriev, L.V., 1989. Noble gases in basalt from a mid-Atlantic Ridge topographic high at 14 degrees N: Geodynamicconsiderations. Earth and Planetary Science Letters 96, 119–133.

Staudigel, H., Zindler, A., Hart, S.R., Leslie, T., Chen, C.T.,Clague, D., 1984. The isotope systematics of a juvenle in-traplate volcano: Pb, Nd, and Sr isotope ratios of basalts fromLoihi seamount, Hawaii. Earth and Planetary Science Letters69, 13–29.

Staudigel, H., Park, K.H., Pringle, M., Rubenstone, J.M., Smith,W.H., Zindler, A., 1991. The longevity of the South Pacificisotopic and thermal anomaly. Earth and Planetary ScienceLetters 102, 24–44.

Stein, M., Hoffman, A.W., 1994. Mantle plumes and episodiccrustal growth. Nature 372, 63–68.

Stein, C.A., Stein, S., 1993. Constraints on Pacific mid-plateswells from global depth-age and and heat-flow age models, inThe Mesozoic Pacific: geology, tectonics and volcanism.American Geophysical Union, Geophysical Monograph 77,pp. 53–76.

Stein, J.S., Fisher, A.T., Sangseth, M., Jin, W., Iturrino, G., Davis,E., 1998. Fine-scale heat flow, shallow heat sources anddecoupled circulation systems at two sea floor hydrothermalsites, Middle Valley, northern Juan de Fuca Ridge. Geology26, 1115–1118.

Steinitz, G., Kapusta, Y., Sandler, A., Kotlarsky, P., 1996. Sedi-mentary K–Ar signatures in clay fractions from marine shelfenvironments. Sedimentology 42, 921–934.

Stewart, C.A., Rampino, M.R., Robinson, C., 1993. Impact shocksin the transition zone: enough energy to trigger a plume?w x Ž .abs. . Eos Transactions, American Geophysical Union 74Ž .43 , 80, Supplement.

Stille, P., Unruh, D.M., Tatsumoto, M., 1986. Pb, Sr, and Hfisotopic constraints on the origin of Hawaiian basalts andevidence for a unique mantle source. Geochimica et Cos-mochimica Acta 50, 2303–2319.

Stuart, G.W., Zengini, T.G., 1987. Seismic crustal structure of theLimpopo mobile belt, Zimbabwe. Tectonophysics 144, 323–335.

Stolper, E.M., Asimow, P.D., Hirschmann, M.M., 1996. Adiabaticw xmelting of the mantle abs . Geological Society of America,

Ž .Abstracts with Programs 28 7 , 290.Ž .Storey, B.C., Alabaster, T., Pankhurst, R.J. Eds. , 1992. Magma-

tism and the Causes of Continental Breakup. Geological Soci-ety, London, Special Publication, vol. 68.

Su, W.J., Woodward, R.L., Dziewonski, A.M., 1992. Deep originof mid-ocean ridge seismic velocity anomalies. Nature 360,149–152.

Sun, Y., Puttmann, W., Speczik, S., 1995. Differences in thedepositional environment of basal Zechstein in southwestPoland. Organic Geochemistry 23, 819–835.

Sundvik, M., Larson, R.L., Detrick, R.S., 1984. Rough–smoothbasement boundary in the western North Atlantic. Geology 12,31–34.

Suwa, K., Oana, S., Wada, H., Osaki, S., 1975. Isotope geochem-istry and petrology of African carbonatites. Physics and Chem-istry of the Earth 9, 435–445.

Sykes, L.R., 1967. Mechanism of earthquakes and nature offaulting on mid-ocean ridges. Journal of Geophysical Research72, 2131–2153.

Tackley, P.J., 1996. Effects of strongly variable viscosity onthree-dimensional compressible convection in planetary man-tles. Journal of Geophysical Research 101, 3311–3332.

Takahashi, E., Kushiro, I., 1983. Melting of dry peridotite at highpressures and basalt magma genesis. American Mineralogist68, 859–879.

Takahashi, E., Nakajima, K., Wright, T.L., 1998. Origin of theColumbia River basalts: melting model of a heterogeneousplume head. Earth and Planetary Science Letters 162, 63–80.

Talwani, M., Windisch, C.C., Langseth, M.G., 1971. ReykjanesRidge crest: a detailed geophysical study. Journal of Geophys-ical Research 76, 473–517.

Tanaka, T., Lin, J., Tucholke, B.E., Klinerock, M.C., Isezaki, N.,1996. Magnetic structure of slow-spreading crustal segmentsinferred from magnetic vectors on the Mid-Atlantic Ridgew x Ž .abs . Eos Transactions, American Geophysical Union 77Ž .46 , 728, Supplement.


Tanimoto, T., Anderson, D.L., 1984. Mapping convection in themantle. Geophysical Research Letters 11, 287–290.

Ž .Tankard, A.J., Balkwill, H.R. Eds. , Extensional Tectonics andStratigraphy of North Atlantic Margins. American Associationof Petroleum Geologists, Memoir, vol. 46.

Tankard, A.J., Jackson, M.P.A., Eriksson, K.A., Hobday, D.K.,Ž .Hunter, D.R., Minter, W.E.L. Eds. , Crustal Evolution of

Southern Africa: 3.8 Billion years of Earth History. Springer-Verlag.

Tarduno, J.A. et al., 1991. Rapid formation of the Ontong–JavaPlateau by Aptian mantle plume volcanism. Science 254,399–403.

Tatsumi, Y., Shinjoe, H., Ishizuka, H., Sager, W.W., Klaus, A.,1998. Geochemical evidence for a mid-Cretaceous super-plume. Geology 26, 151–154.

Tilley, C.E., 1952. Some trends of basaltic magma in limestonesyntexis. American Journal of Science, Bowen, 529–545.

Tivey, M.A., Tucholke, B.E., 1998. Magentization of 0–29 Maocean crust on the Mid-Atlantic Ridge. Journal of GeophysicalResearch 103, 17807–17826.

Trumpy, R., 1975. Penninic–Austroalpine boundary in the SwissAlps: a presumed former continental margin and its problems.American Journal of Science 275A, 209–238.

Toft, P.B., Arkani-Hamed, J., 1992. Magnetization of the PacificOcean lithosphere deduced from magsat data. Journal of Geo-physical Research 97, 4387–4406.

Toomey, D.R., Wilcock, W.S., Solomon, S.C., Hammond, W.C.,Orcutt, J.A., 1998. Mantle seismic structure beneath the MELTregion of the East Pacific Rise from P and S wave tomogra-phy. Science 280, 1224–1227.

Tucholke, B.E., Stewart, W.K., Kleinrock, M.C., 1997. Long-termdenudation of ocean crust in the central North Atlantic. Geol-ogy 25, 171–174.

Turcotte, D.L., 1982. Driving mechanism of mountain building.Ž .In: Hsu, K.J. Ed. , Mountain Building Processes. Academic

Press, London, pp. 141–146, 262 pp.Turcotte, D.L., Schubert, G., 1982. Geodynamics. Wiley, New

York.Turner, D.L., Jarrard, R.D., 1982. K–Ar dating of the Cook–

Austral chain: a test of the hot spot hypothesis. Journal ofVolcanology and Geothermal Research 12, 187–220.

Tuttle, O.F., Bowen, N.L., 1958. Origin of granite in the light ofexperimental studies. Geological Society of America, Memoir,74.

Ulmer, P., Tromsdorff, V., 1995. Serpentine stability to mantledepth and subduction-related magmatism. Science 268, 858–861.

Valbracht, P.J., Honda, M., Staudigel, H., McDougall, I., Trost,A.P., 1994. Noble gas partitioning in natural samples: resultsfrom coexisting glass and olivine phenocrysts in four Hawai-

Ž .ian submarine basalts. In: Matsuda, J. Ed. , Noble Gas Geo-chemistry and Cosmochemistry. Terra Scientific Publishing,Tokyo, pp. 373–381.

Valbracht, P.J., Staudigel, H., Honda, M., McDougall, I., Davies,G.R., 1996. Isotopic tracing of volcanic source regions fromHawaii: decoupling of gaseous from lithophile magma compo-nents. Earth and Planetary Science Letters 144, 185–198.

Valbracht, P.J., Staudacher, T., Malahoff, A., Allegre, C.J., 1997.Noble gas systematics of deep rift zone glasses from Loihiseamount, Hawaii. Earth and Planetary Science Letters 150,399–411.

Van Andel, Tj.H., 1994. New Views of an Old Planet, 2nd edn.Cambridge Univ. Press.

Van der Hilst, R.D., Widiyantoro, S., Engdahl, E.R., 1997. Evi-dence for deep mantle circulation from global tomography.Nature 386, 578–584.

Van der Merwe, S.W., 1995. The relationship between thrusting,vertical shears and open folds in the western part of theNamaqua mobile belt, South Africa. South African Journal ofGeology 98, 68–77.

Vasco, D.W., Johnson, L.R., Pulliam, R.J., Earle, P.S., 1994.Robust inversion of travel-time residuals for mantle P- andS-velocity structure. Journal of Geophysical Research 99,13727–13755.

Vauchez, A., Nicolas, A., 1991. Mountain building: strike-parallelmotion and mantle anisotropy. Tectonophysics 185, 183–201.

Vauchez, A., Barrual, G., Nicolas, A., 1999. Comment on ASKSsplitting beneath continental rift zonesB by Gao et al. Journalof Geophysical Research 104, 10787–10789.

Veizer, J., 1989. Strontium isotopes in seawater through time.Annual Reviews of Earth and Planetary Sciences 17, 141–167.

Veizer, J., Bell, K., Jansen, S.L., 1992. Temporal distribution ofcarbonatites. Geology 20, 1147–1149.

Verosub, K.L., Moores, E.M., 1981. Tectonic rotations in exten-sional regimes and their paleomagnetic consequences foroceanic basalts. Journal of Geophysical Research 86, 6335–6349.

Vidal, P., 1992. More HIMU in the future? Geochimica et Cos-mochimica Acta 52, 4295–4299.

Vidal, P., Chavel, C., Brousse, R., 1984. Large mantle heterogene-ity beneath French Polynesia. Nature 307, 536–538.

Vinnik, L.P., Makeyeva, L.I., Milev, A., Usenko, Y., 1992.Global pattern of azimuthal anisotropy and deformation in thecontinental mantle. Geophysical Journal International 111,433–447.

Vinnik, L.P., Krishna, V.G., Kind, R., Bormann, P., Stammler,K., 1994. Shear wave splitting in the records of the Germanregional seismic network. Geophysical Research Letters 21,457–460.

Vogt, P.R., 1971. Asthenospheric motion recorded by ocean floorsouth of Iceland. Earth and Planetary Science Letters 13,153–160.

Ž .Vogt, P.R., Tucholke, B.E. Eds. , The Western North AtlanticRegion. Geological Society of America, Geology of North

wAmerica, vol. M see chapters by Einarsson, Kent and Grad-stein, Macdonald, Melson and OHearn, Palmason, Saemunds-

xson, Schilling, Vogt .Von Herzen, R.P., Cordery, M.J., Detrick, R.S., Fang, C., 1989.

Heat flow and the thermal origin of hotspot swells: the Hawai-ian swell revisited. Journal of Geophysical Research 94,13783–13799.

Von Herzen, R.P., 1991. Proceedings, Ocean Drilling Program,Scientific Results, Leg 118: College Station, pp. 541–556.

Walker, G.P., 1960. Zeolite zones and dyke distribution in relation


to the structure of basalts in eastern Iceland. Journal ofGeology 68, 515–528.

Walker, G.P., 1975. Intrusive sheet swarms and the identity ofcrustal layer 3 in Iceland. Journal of the Geological Society,London 131, 143–161.

Watts, A.B., ten Brink, U.S., 1989. Crustal structure, flexure andsubsidence history of the Hawaiian islands. Journal of Geo-physical Research 94, 10473–10500.

Wetherill, G.W., 1954. Variations in the isotopic abundances ofneon and argon extracted from radioactive minerals. PhysicsReviews 96, 679–683.

Weaver, B.L., 1991. The origin of oceanic island basalt end-mem-Žbers. Earth and Planetary Science Letters 104, 381 cf. Geol-

.ogy 19, 123 .Weinstein, S.A., 1993. Catastrophic overturn of the Earth’s mantle

driven by multiple phase changes and internal heat generation.Geophysical Research Letters 20, 101–104.

Weis, D., Frey, F.A., 1991. Isotope geochemistry of NinetyeastŽ .Ridge basalts. In: Weissel, D. Ed. , 1991. Proceedings of the

Ocean Drilling Program, Scientific Results, vol. 121, pp.w591–610 see related papers by Resiwati, Duncan, Smith et al.,

xSaundes et al. .Weissel, J.K., Pierce, J. et al., 1991. Proceedings Ocean Drilling

Project, Scientific results, vol. 121, College Station, TX.Wessel, P., 1993. A re-examination the flexural deformation

beneath the Hawaiian islands. Journal of Geophysical Re-search 98, 12177–12190.

White, W.M., 1985. Sources of oceanic basalt: radiogenic isotopicevidence. Geology 13, 115–118.

White, R.S., Bown, J.W., Smallwood, J.R., 1995. The temperatureof the Iceland plume and the origin of outward-propagatingV-shaped ridges. Journal of the Geological Society, London152, 1039–1045.

White, W.M., Patchett, P.J., Othman, D.B., 1984. U, Th and Pb inmarine sediments: crustal recycling and the isotopic evolutionof mantle lead. EOS 65, 296.

Wickman, F.E., 1948. Isotope ratios: a clue to the age of certainmarine sediments. Journal of Geology 56, 61–66.

Widom, E., Carlson, R.W., Gill, J.B., Schmincke, H.U., 1997.Th–Sr–Nd–Pb isotope and trace element evidence for theorigin of the Sao Miguel, Azores, enriched mantle source.Chemical Geology 140, 49–68.

Wilcox, W.S., Dougherty, M.E., Solomon, S.C., Purdy, G.M.,Toomey, D.R., 1993. Seismic propagation across the EastPacific Rise: implications for seismic tomography. Journal ofGeophysical Research 98, 19913–19932.

Willet, S.D., Beaumont, C., 1994. Subduction of Asian litho-spheric mantle beneath Tibet inferred from models of conti-nental collision. Nature 369, 642–645.

Wilson, J.T., 1963. A possible origin of the Hawaiian Islands.Canadian Journal of Physics 41, 863–870.

Winterer, E.L., 1998. Cretaceous karst guyots: new evidence forsubaerial erosional terrain. Geology 26, 59–62.

Wolfe, C., Bjarnason, I., Vandecar, J., Solomon, S., 1997. Seis-mic structure of the Iceland mantle plume. Nature 385, 245–247.

Woodhead, J.D., Devey, C.W., 1993. Geochemistry of the Pitcairnseamounts. Earth and Planetary Science Letters 116, 81–99.

Woodhouse, J.H., Dziewonski, A.M., 1989. Seismic modeling ofthe Earth’s large scale 3-dimensional structure. PhilosophicalTransactions of the Royal Society of London, Ser. A 328,291–308.

Woods, M.T., Leveque, J.J., Okai, E.A., Cara, N., 1991. Two-sta-tion measurement of Rayleigh wave group velocity along theHawaiian swell. Geophysical Research Letters 18, 105–108.

Woodward, R.L., Schnieder, R., 1993. High resolution surfacew x Žwave tomography of North America abs. . Eos Transactions,

. Ž .American Geophysical Union 74 16 , 203, Supplement.Wooley, A.R., 1989. The spacial and temporal distribution of

Ž .carbonatites. In: Bell, K. Ed. , Carbonatites: Genesis andEvolution. Unwin-Hyman, London, pp. 15–37.

Worzel, J.L., 1965. Deep structure of the continental margins andŽ .mid-ocean ridges. In: Whittard, W.F., Bradshaw, R. Eds. ,

Submarine Geology and Geophysics. Butterworths, London,pp. 35–361.

Wyllie, P.J., Wolf, M.B., 1993. Amphibole dehydration melting:sorting out the solidus. In: Pritchard, H.M., Alabaster, T.,

Ž .Harris, N.B., Neary, C.R. Eds. , Magmatic Processes andPlate Tectonics. Geological Society, London, Special Publica-

Žtion, vol. 76, pp. 405–416 cf. sequel in Journal of Geophysi-.cal Research 100, 15601–15621 .

Yingping, L., Thurber, C.H., Munson, C.G., 1992. Profile ofdiscontinuities beneath Hawaii from S to P converted seismicwaves. Geophysical Research Letters 19, 111–114.

Yoder, H.S., 1976. Generation of Basaltic Magma. NationalAcademy of Sciences, Washington, DC.

Ž .Yoder, H.S. Ed. , Evolution of the Igneous Rocks: Fiftieth An-niversary Perspectives. Princeton Univ. Press.

Zack, T., Foley, S.F., Rivers, T., 1999. The role of phengite,zoisite, lawsonite and rutile as trace element carriers in sub-

w x Žduction slabs abs. . Eos Transactions, American Geophysical. Ž .Union 80 17 , S350.

Zartman, R.E., Haines, S.M., 1988. The plumbotectonic model forPb isotope systematics among major terrestrial reservoirs.Geochimica et Cosmochimica Acta 52, 1327–1339.

Zhang, S., Karato, S.I., 1995. Lattice preferred orientation ofolivine aggregates deformed in simple shear. Nature 375,774–777.

Zhang, Y.S., Tanimoto, T., 1993. High resolution global uppermantle structure and plate tectonics. Journal of GeophysicalResearch 98, 9793–9823.

Zhang, Y.S., Tanimoto, T., Stolper, E.M., 1994. S-wave velocity,basalt chemistry and bathymetry along the Mid-Atlantic Ridge.

ŽPhysics of Earth and Planetary Interiors 84, 79–93 cf. Nature.355, 45–49 .

Zhou, R., Grand, S.P., Tajima, F., 1995. High velocity zonew x Žbeneath the southern Tibetan Plateau abs. . Eos Transactions,

. Ž .American Geophysical Union 76 46 , 402, Supplement.Zindler, A., Hart, S.R., 1986. Chemical geodynamics. Annual

Review of Earth and Planetary Sciences 14, 493–571.Zindler, A., Bourdon, B., Elliott, T., 1994. Debunking the k

conundrum. United States Geological Survey, Circular 1107,371.

Zverev, S.M., Litvinenko, I.V., Palmason, G., Yaroshevskaya,G.A., Osokin, M.N., 1980. A seismic study of the axial riftzone in Iceland. Journal of Geophysics 47, 191–210.


Dr. Keith studied geological sciences at University of Alberta,Queens University in Kingston, Ontario, and at MIT. He taughtpetrology and geophysics at Queens, and after 4 years of experi-mental petrology at the Geophysical Laboratory in Washington,DC, moved to Penn State University in 1950. His field work wasin the Canadian shield, and included exploration and mapping forthe Geological Survey of Canada, Ontario Department of Minesand several mining companies. He taught several courses ingeochemistry at Penn State, and is currently Emeritus Professor ofGeochemistry. He has spent recent years reviewing the evidencefor and against the popular plate tectonics model.

Documents

Evidence for a plate tectonics debate