and conservative estimates can be made for their geometry, orientation, and location based on available geophysical and geological data while adhering to the rules of plate
tectonics. Pre-Jurassic ocean basin physiographies have
never been explicitly reconstructed to date.
Preserved magnetic lineations in the Pacific Ocean provide unequivocal evidence that a vast mid-ocean ridge
system existed in the Pacific Ocean in the mid/late Cretaceous, significantly longer than today’s ridge system, and
much of which is now subducted – this was first recognized by Larson and Chase (1972). Subsequent detailed
mapping and compilation of magnetic M-sequence anomalies in the northwestern Pacific Ocean (Nakanishi et al.,
1992a) revealed the complete Mesozoic magnetic anomaly lineation pattern in this area and the origin of the
Pacific plate as a triangularly shaped microplate at a triple
junction between the Izanagi, Farallon, and Phoenix plates
in the Panthalassa Ocean basin, the predecessor of the
Pacific Ocean. This work forms the basis of Pacific Ocean
reconstructions. In the southwest Pacific, the reconstructions from Seton et al. (2012) implement the opening of
ocean basins between fragments of the Ontong Java–
Manihiki–Hikurangi large igneous provinces (LIPs)
(Viso et al., 2005; Taylor, 2006), a major tectonic event
that was missing from previous reconstructions of the Cretaceous Pacific Ocean (Müller et al., 2008b). The Tethys
Ocean, separating Laurasia from Gondwanaland, is
reconstructed by accounting for the migration history of
a series of Gondwanaland ribbon continents that were
rifted off the northern margin of Gondwanaland and
accreted to southern Eurasia in several stages. By using
combined evidence from preserved magnetic lineations
in the present-day Indian Ocean and geological data from
the terranes that accreted to southern Eurasia and the rules
of plate tectonics (Cox and Hart, 1986), it is possible to
constrain the overall geometries of the Tethyan
mid-ocean ridge system. The closure of the Mongol–
Okhotsk Ocean between 200 and 150 Ma is modeled after
van der Voo et al. (1999).
The overall physiography of the global oceans since the
Mesozoic follows an alternating pattern mid-ocean ridge
length increase and decrease as well as changes in the
average age of the oceanic lithosphere. A number of old
mid-ocean ridge flanks in Panthalassa and the Tethys
oceans were gradually being destroyed between 200 and
150 Ma, while new mid-ocean ridge systems were initiated in a stepwise fashion between the late Jurassic and
early Cretaceous, leading to a substantial younging and
shallowing of the ocean basins from the Jurassic to the
Cretaceous period. After the Cretaceous, the length of
the mid-ocean ridge system decreased and was accompanied by an aging and deepening of the ocean basins toward
the present day.
To produce a model of the age-area distribution of the
ocean basins since the Mesozoic, we take the global set
of plate rotations, reconstructed plate boundaries and seafloor magnetic anomalies, and fracture zones imbedded in
the global plate model of Seton et al. (2012) and construct
a set of seafloor spreading isochrons, equivalent to paleomid-ocean ridge geometries and locations. Seafloor isochrons are then gridded (Müller et al., 2008a) to produce
smooth, continuous models for the age-area distribution
of ocean floor through time (see “Plate Motion”). The
resultant age grid becomes the primary input to producing
models of the depth of the ocean basins through time,
based on a selected depth–age relationship. To convert
age to depth, we use the GDH1 model (Stein and Stein,
1992). Other recent depth–age models, e.g., Crosby and
McKenzie (2009), result in an extremely similar mapping
of age to depth (see Müller et al. (2008a) for details) and
would make little difference for the reconstruction of
ocean basin depths.
Oceanic plateaus and sediment cover
Two additional factors play a significant role in controlling ocean basin depth through time, namely, the generation of oceanic large igneous provinces (LIPs) and
oceanic sedimentation. To account for the effect of LIPs
on oceanic depth, we take a subset of the Coffin
et al. (2006) dataset to include LIPs that were generated
by a transient plume head as we are most interested in capturing the voluminous volcanism rather than the small
scale. We use Schubert and Sandwell’s (1989) method to
determine the average elevation of major oceanic plateaus
relative to the surrounding seafloor, based on the difference between the modal depths within two polygons,
one outlining the perimeter of a given oceanic plateau
and another larger polygon including surrounding ocean
floor (Figure 1). Individual plateaus are included in basement depth grids at their emplacement time based on a
compilation of eruption ages (Müller et al., 2008b). Where
available, the paleo-depth estimates of LIPs were crossreferenced with published data (e.g., back-stripping from
well data or from seismic data).
As the ocean floor ages, the thickness of its sediment
cover increases, but abyssal sediment thickness is also latitude dependent, as illustrated in a polynomial surface fit
of global sediment thickness as a function of oceanic
crustal age and latitude (Figure 2). The observed latitudinal sediment thickness variation primarily reflects differences in productivity between zones of tropical
upwelling and an increase in sediment thickness toward
high latitudes, reflecting terrigenous runoff. The simple
relationship provides a good first approximation for
reconstructing an estimate for regional oceanic sediment
thickness through time. We isostatically correct for sediment loading based on Sykes (1996).
Paleophysiography of Mesozoic and Cenozoic
ocean basins
Overview
Broad-scale paleo-bathymetries of the ocean basins can be
constructed by combining oceanic basement depths
derived from a depth–age relationship with oceanic plateaus and sediment thickness estimates (Figure 3). The
PALEOPHYSIOGRAPHY OF OCEAN BASINS
639
tectonics. Pre-Jurassic ocean basin physiographies have
never been explicitly reconstructed to date.
Preserved magnetic lineations in the Pacific Ocean provide unequivocal evidence that a vast mid-ocean ridge
system existed in the Pacific Ocean in the mid/late Cretaceous, significantly longer than today’s ridge system, and
much of which is now subducted – this was first recognized by Larson and Chase (1972). Subsequent detailed
mapping and compilation of magnetic M-sequence anomalies in the northwestern Pacific Ocean (Nakanishi et al.,
1992a) revealed the complete Mesozoic magnetic anomaly lineation pattern in this area and the origin of the
Pacific plate as a triangularly shaped microplate at a triple
junction between the Izanagi, Farallon, and Phoenix plates
in the Panthalassa Ocean basin, the predecessor of the
Pacific Ocean. This work forms the basis of Pacific Ocean
reconstructions. In the southwest Pacific, the reconstructions from Seton et al. (2012) implement the opening of
ocean basins between fragments of the Ontong Java–
Manihiki–Hikurangi large igneous provinces (LIPs)
(Viso et al., 2005; Taylor, 2006), a major tectonic event
that was missing from previous reconstructions of the Cretaceous Pacific Ocean (Müller et al., 2008b). The Tethys
Ocean, separating Laurasia from Gondwanaland, is
reconstructed by accounting for the migration history of
a series of Gondwanaland ribbon continents that were
rifted off the northern margin of Gondwanaland and
accreted to southern Eurasia in several stages. By using
combined evidence from preserved magnetic lineations
in the present-day Indian Ocean and geological data from
the terranes that accreted to southern Eurasia and the rules
of plate tectonics (Cox and Hart, 1986), it is possible to
constrain the overall geometries of the Tethyan
mid-ocean ridge system. The closure of the Mongol–
Okhotsk Ocean between 200 and 150 Ma is modeled after
van der Voo et al. (1999).
The overall physiography of the global oceans since the
Mesozoic follows an alternating pattern mid-ocean ridge
length increase and decrease as well as changes in the
average age of the oceanic lithosphere. A number of old
mid-ocean ridge flanks in Panthalassa and the Tethys
oceans were gradually being destroyed between 200 and
150 Ma, while new mid-ocean ridge systems were initiated in a stepwise fashion between the late Jurassic and
early Cretaceous, leading to a substantial younging and
shallowing of the ocean basins from the Jurassic to the
Cretaceous period. After the Cretaceous, the length of
the mid-ocean ridge system decreased and was accompanied by an aging and deepening of the ocean basins toward
the present day.
To produce a model of the age-area distribution of the
ocean basins since the Mesozoic, we take the global set
of plate rotations, reconstructed plate boundaries and seafloor magnetic anomalies, and fracture zones imbedded in
the global plate model of Seton et al. (2012) and construct
a set of seafloor spreading isochrons, equivalent to paleomid-ocean ridge geometries and locations. Seafloor isochrons are then gridded (Müller et al., 2008a) to produce
smooth, continuous models for the age-area distribution
of ocean floor through time (see “Plate Motion”). The
resultant age grid becomes the primary input to producing
models of the depth of the ocean basins through time,
based on a selected depth–age relationship. To convert
age to depth, we use the GDH1 model (Stein and Stein,
1992). Other recent depth–age models, e.g., Crosby and
McKenzie (2009), result in an extremely similar mapping
of age to depth (see Müller et al. (2008a) for details) and
would make little difference for the reconstruction of
ocean basin depths.
Oceanic plateaus and sediment cover
Two additional factors play a significant role in controlling ocean basin depth through time, namely, the generation of oceanic large igneous provinces (LIPs) and
oceanic sedimentation. To account for the effect of LIPs
on oceanic depth, we take a subset of the Coffin
et al. (2006) dataset to include LIPs that were generated
by a transient plume head as we are most interested in capturing the voluminous volcanism rather than the small
scale. We use Schubert and Sandwell’s (1989) method to
determine the average elevation of major oceanic plateaus
relative to the surrounding seafloor, based on the difference between the modal depths within two polygons,
one outlining the perimeter of a given oceanic plateau
and another larger polygon including surrounding ocean
floor (Figure 1). Individual plateaus are included in basement depth grids at their emplacement time based on a
compilation of eruption ages (Müller et al., 2008b). Where
available, the paleo-depth estimates of LIPs were crossreferenced with published data (e.g., back-stripping from
well data or from seismic data).
As the ocean floor ages, the thickness of its sediment
cover increases, but abyssal sediment thickness is also latitude dependent, as illustrated in a polynomial surface fit
of global sediment thickness as a function of oceanic
crustal age and latitude (Figure 2). The observed latitudinal sediment thickness variation primarily reflects differences in productivity between zones of tropical
upwelling and an increase in sediment thickness toward
high latitudes, reflecting terrigenous runoff. The simple
relationship provides a good first approximation for
reconstructing an estimate for regional oceanic sediment
thickness through time. We isostatically correct for sediment loading based on Sykes (1996).
Paleophysiography of Mesozoic and Cenozoic
ocean basins
Overview
Broad-scale paleo-bathymetries of the ocean basins can be
constructed by combining oceanic basement depths
derived from a depth–age relationship with oceanic plateaus and sediment thickness estimates (Figure 3). The
PALEOPHYSIOGRAPHY OF OCEAN BASINS
639