Observational studies on the temporal and spatial
structures of the deep-circulation current in thewestern North Pacific

西部北太平洋における深層循環流の時空間的構造の
観測的研究

柳本

大吾

Observational studies on the temporal and spatial
structures of the deep-circulation current in thewestern North Pacific

A dissertation submitted for the degree of Doctor of Science

BY

Daigo Yanagimoto
Atmosphere and Ocean Research Institute, The University of Tokyo

Abstract
For understanding the global deep ocean circulation, which occupies large part of the global
meridional overturning circulation, knowing its pathway and volume transport is essential. The
author clarified the pathway and the volume transport variability of the deep-circulation current
in the western North Pacific with analyzing current velocity data directly measured at two
mooring arrays along 165°E and southwest of the Shatsky Rise. The both arrays were deployedat key points for the deep-circulation current in the region between 20°N and 35°N in the
Northwest Pacific Basin. Five moorings were deployed at 27°–35°N along 165°E from 1991 to
1993 and additional one mooring at 27°N, 168°E from 1989 to 1991. Nine moorings were
deployed between the Shatsky Rise and the Ogasawara Plateau (an eastward bulge from the
Izu-Ogasawara Ridge near 25°N) from 2004 to 2005 and additional two moorings on the foot of
the Shatsky Rise from 2003 to 2004. The author also utilized recent high-quality andhigh-resolution dissolved-oxygen data in order to estimate the pathway of the deep-circulation
current passing the moorings.
First, the pathway of the deep-circulation current was examined. The eastern branch of the
deep-circulation current passes north-northeastward through the Wake Island Passage, turns
cyclonically with flowing northwestward around 33°N, 165°E, takes a detour south of the
Shatsky Rise, then, flows northwestward mainly on the lower part of the slope of the Shatsky
Rise. The deep-circulation current is not found to flow eastward between the Wake-Necker
Ridge and the Hess Rise into the Northeast Pacific Basin. The western branch of the
deep-circulation current proceeds northward beyond 25°N with a detour east of the Ogasawara
Plateau. Thus, both of the eastern and western branches enter the central region of theNorthwest Pacific Basin through the relatively narrow strait between the Shatsky Rise and the
Ogasawara Plateau.
Next, the temporal variations of volume transports of the western and eastern branches were
examined southwest of the Shatsky Rise. Almost zero volume transport occurs at intervals of
2–4 months in the both branches. They have dominant spectral peaks at periods of 3 months and1month. This is similar to the volume transport variation at the Wake Island Passage in the
North Pacific, while it is not similar to those at the Samoan Passage and in the Southwest
Pacific Basin in the South Pacific.
Additionally, volume transport variation of the eastern branch has high correlation with its
width variation. The eastern branch flowing with an average width of 200 ± 70 km
intermittently expands up to 200 km away southwestward from its main position on the loweripart of the slope of the Shatsky Rise. The western branch is so narrow that we find it almost
exclusively at one mooring system northeast of the Ogasawara Plateau.
These findings do not only elucidate the pathway of the deep circulation in the western
North Pacific that has been different for researchers so far, but also provide an importantsuggestion on the northward extension of the meridional overturning circulation in the Pacific to
the Aleutian Trench.

iiAcknowledgement
I would like to express my sincere thanks to Prof. M. Kawabe for his invaluable suggestions
and advices and continuous encouragement. I would like to express my great thanks to Prof. S.
Fujio for his invaluable suggestions and advices. I greatly thank to Drs. E. Oka and K. Tanaka
for their helpful suggestions and advices and to Mr. S. Kitagawa for his excellent technicalworks. I greatly thank to Profs. H. Hasumi, T. Hibiya, I. Yasuda, M. Sato, and K. Iga for theirhelpful suggestions. I would like to express my thanks to my colleagues in the former Ocean
Research Institute and the existing Atmosphere and Ocean Research Institute for fruitful
discussions. I would like to express my hearty thanks to Prof. Emeritus K. Taira for his
continuous encouragement.
The mooring works, that the present thesis is essentially based on, were carried out incruises KH-91-5, KH-93-2, KH-03-1, KH-04-4, and KH-05-4 of the R/V Hakuho-maru. The
hydrographic observations, that the present thesis is greatly based on, were carried out in thesecruises as well as the additional ones of the R/V Hakuho-maru, i.e., KH-95-1, KH-99-1,
KH-07-1. I am greatly indebted to the captains and crews of the R/V Hakuho-maru, and thescientists who joined the cruises with the author.
I am greatly thankful to my family.

iii

Contents
Chapter 1. General introduction ................................................................................. 1
Chapter 2. Temporal and spatial characteristics of deep-circulation current
measured at mid-latitude of 165°E ....................................................................... 13
2.1.

Introduction .................................................................................................................. 13

2.2.

Mooring observations .................................................................................................. 14

2.3.

Temporal and spatial variations of near-bottom currents at 165°E .............................. 14

2.4.

Frequency-domain characteristics of current variations .............................................. 17

2.5.

Conclusion ................................................................................................................... 18

Chapter 3. Pathways of the deep-circulation currents in the southern part of theNorthwest Pacific Basin ......................................................................................... 28
3.1.

Introduction .................................................................................................................. 28

3.2.

Data and method .......................................................................................................... 28

3.3.

Pathways of the deep-circulation current and dissolved oxygen distributions ............ 29

3.4.

Conclusion ................................................................................................................... 30

Chapter 4. Temporal and spatial characteristics of current velocities andvolume transports of deep-circulation currents measured southwest of the
Shatsky Rise ............................................................................................................ 39
4.1.

Introduction .................................................................................................................. 39

4.2.

Mooring observations .................................................................................................. 40

4.3.

Velocity field near the ocean bottom............................................................................ 41

4.4.

Variability in volume transport of the deep circulation ................................................ 43

4.5.

Conclusion ................................................................................................................... 44

Chapter 5. General conclusion .................................................................................. 61
Appendix A. Dependency of volume transport estimation from current meterdata on interpolation method ................................................................................ 66
Appendix B. Anticyclonic eddy over an isolated seamount ...................................... 71
References...................................................................................................................... 75

iv

Chapter 1. General introduction
In the ocean, there are horizontal wind-driven circulations as well as a global meridional
overturning circulation (MOC), which is associated with sinking at high latitudes in the North
Atlantic and the Southern Ocean and upwelling elsewhere. The MOC takes a significant role
for moderating the world climate by transporting cold waters from high latitudes to low
latitudes in the lower layer and warm waters from low latitudes to high latitudes in the upper
layer. Therefore, the MOC is attracting attention of oceanographers in recent years in particular
as is reviewed by Richardson (2007). The deep circulation occupies the thick deep portion of
the MOC and contributes greatly to transporting heat and materials. The MOC in the Pacific
Ocean, where there is no significant sinking of water from the surface to the abyssal layer,
consists of the deep circulation flowing northward from the Southern Ocean and southwardfrom the North Pacific. It is a large question how far north the MOC extends in the Pacific
Ocean.
The global deep circulation originates from formation of saline, oxygen-rich deep water,
North Atlantic Deep Water (NADW), mainly off Greenland. NADW is transported southward
to the Southern Ocean, where NADW is reformed to Lower Circumpolar Deep Water (LCDW)
with affected by cold Antarctic Bottom Water (AABW) formed around Antarctica. Thus,
LCDW does not only have low temperature but also inherits a feature of high salinity and highdissolved oxygen from NADW. Such dense, oxygen-rich LCDW spreads along the bottom toeach ocean from the Southern Ocean with transported northward by the deep circulation.
Therefore, high dissolved oxygen and low silica of LCDW is used as a tracer of the deep
circulation (e.g., Mantyla and Reid, 1983). The main stream of the deep circulationtransporting LCDW is referred as “deep-circulation current” in this thesis.
The South Pacific has two entrances for LCDW, i.e., the Southwest Pacific Basin and theChile Basin (Mantyla and Reid, 1983) (Fig. 1.1). While LCDW entering the Chile Basin does
not proceed northward beyond the equator due to sill depth, LCDW entering the Southwest
Pacific Basin is transported northward further to the North Pacific. Because the Southwest
Pacific Basin has simple bottom topography, LCDW and the deep-circulation current are foundalong the Tonga-Kermadec Ridge, the western boundary of the basin (Fig. 1.2), as is expectedfrom the idealized-model theory by Stommel and Arons (1960). At 32°30′S in the Southwest
Pacific Basin, LCDW is found as a high-salinity core of 34.72–34.73 PSU below 3000 m depth,
and it has high dissolved oxygen (4.8 ml l-1) and low temperature (0.58°C) (Whitworth et al.,1999). The volume transport of LCDW at this latitude is estimated at 12 Sv (1 Sv = 106 m3 s-1)
as a geostrophic transport based on inversely-solved reference velocities by Tsimplis et al.

1

(1998).
After flowing northward along the Tonga-Kermadec Ridge in the Southwest Pacific Basin,
LCDW reaches series of ridges from the Manihiki Plateau to the Robby Ridge (Fig. 1.2).
LCDW passes through the Samoan Passage, the eastern flank of the Manihiki Plateau, and agap of the Robby Ridge, and enters the Central Pacific Basin. Roemmich et al. (1996) estimate
volume transports of LCDW (θ<1.1°C) passing these passages at 7.8 Sv in the Samoan
Passage, 2.8 Sv at the eastern flank of the Manihiki Plateau, and 1.1 Sv in the Robby-Ridge
gap; thus, LCDW is transported with volume transport of 11.7 Sv into the Central Basin. Taftet al. (1991) also estimate geostrophic volume transport in the Samoan Passage at 6.0 Sv.
Distributions of water properties show that LCDW spreads from the Central Pacific Basinto the basins in the North Pacific (Mantyla and Reid, 1983). In contrast with the western South
Pacific, it has been difficult to identify its pathway in the western North Pacific due to theextremely complex bottom topography (Fig. 1.2). Nevertheless, accumulation of precise
observation data in the World Ocean Circulation Experiment (WOCE) in the 1990s andpost-WOCE experiments has made clear the detail of the deep-circulation current in thetropical North Pacific. It is clarified that the deep-circulation current bifurcates into two branchcurrents in the Central Pacific Basin due to sill depth of the western boundary of the basin, theGilbert Ridge. The deep part of the deep-circulation current forms the eastern branch, flowing
northward along the Gilbert Ridge toward the Wake Island Passage with a volume transport of
8.1 Sv, whereas the shallow part forms the western branch flowing into the Melanesian andEast Mariana basins beyond the ridge with a volume transport of 2–6 Sv (Johnson and Toole,
1993; Kawabe, 1993; Kawabe and Taira, 1995; Kawabe et al., 2003; Kawabe et al., 2006) (Fig.
1.2).
The eastern branch has another separation of northward and eastward flows in the Central
Pacific Basin (black arrows in Fig. 1.2). The eastward flow enters the Northeast Pacific Basin(Kato and Kawabe, 2009) and the northward flow enters the Northwest Pacific Basin throughthe Wake Island Passage and adjacent gaps in the Mid-Pacific Seamounts (Kawabe et al.,2003). The eastern branch passing the Wake Island Passage northward is examined with using
moorings by Kawabe et al. (2005). They revealed its mean volume transport of 3.6 ± 1.3 Sv.
The total northward volume transport including those at the adjacent gaps is estimated at morethan 4 Sv.
The western branch in the East Mariana Basin also separates into flows toward adjacentbasins (gray arrows in Fig. 1.2). Some part flows into the West Mariana Basin and some partflows into the East Caroline Basin. Kawabe et al. (2003) estimate the volume transport to the2West Mariana Basin as 0.3 Sv and the one to the East Caroline Basin as 1.7 Sv. Siedler et al.
(2004) confirmed these currents with using moorings and estimate total volume transport to thetwo basins at 1 Sv. The remaining part of the western branch proceeds further northward
through gaps in the Magellan Seamounts with a volume transport of 2.1 Sv (Kawabe et al.,2003).
Thus, the deep-circulation current south of 20°N and the Wake Island Passage wasclarified by Kawabe et al. (2003, 2005). However, the eastern branch north of the Wake Island
Passage is not yet confirmed and is shown with broken lines in Kawabe et al. (2003) (Fig. 1.3).
According to geostrophic velocity and dissolved oxygen distribution obtained in sections of
WOCE Hydrographic Programme (WHP) P13C and P13J at 165°E, Kawabe and Taira (1995)
suppose that the eastern branch turns cyclonically in the small basin surrounded by the Shatsky
Rise, the Emperor Seamount Chain, and the Wake–Necker Ridge after passing the Wake Island
Passage. Eastward outflow near the Midway Island toward the Northeast Pacific Basin inferredfrom dissolved-oxygen distribution by Mantyla and Reid (1983) and northward proceeding of
the western branch with a detour east of the Ogasawara Plateau, an eastward bulge of the
Izu-Ogasawara Ridge at 25°N, suggested by Kawabe et al. (2003) are also drawn with brokenlines in Fig. 1.3.
On the other hand, Wijffels et al. (1998) show water-mass distribution and geostrophic
volume transport from their meridional hydrographic section of WHP P10 at about 149°E.
They indicate westward geostrophic transport (7.9 Sv) in the deepest layer north of 25°N; thisis similar with the westward pathway of the eastern branch supposed by Kawabe and Taira(1995) (Fig. 1.4a). However, they illustrate the western branch transporting LCDW blocked bya benthic water-mass front near 25°N and turning eastward with mixing with the upper water(Fig. 1.4b).
Furthermore, Owens and Warren (2001) illustrate the deep circulation with summarizing
previous results of observations and simulations (Fig. 1.5). They adopt an eastward pathway ofthe deep water passing between the Shatsky Rise and the Mid-Pacific Seamounts after
bifurcating from the western branch north of the Ogasawara Plateau. The eastward current
proceeds nearly along 30°N, joins the eastern branch, goes further eastward, and finally flows
out to the Northeast Pacific Basin as inferred by Mantyla and Reid (1983). Therefore, thecyclonic turn of deep current near 35°N, 165°E is illustrated as just a local gyre out of the mainstream of the deep-circulation current. This eastward pathway of the deep-circulation current
contradicts the pathway of the eastern branch supposed by Kawabe and Taira (1995).
Therefore, the deep-circulation current in the southern region between 20°N and 35°N of3the Northwest Pacific Basin must be clarified and confirmed on the basis of observation data.
The points of issue to be solved are (1) whether part of the eastern branch flows eastward
between the Wake-Necker Ridge and the Hess Rise into the Northeast Pacific Basin or not, (2)whether the western branch proceeds northwestward with a detour east of the Ogasawara
Plateau or not, and (3) whether the deep-circulation current flows westward or eastward
between 20°N and 35°N. These questions (1), (2), and (3) are examined in this thesis.
The answer to these questions will provide an important suggestion particularly on thenorthward extension of the MOC in the Pacific. If the deep circulation passes eastward theregion at about 30°N between the Wake-Necker Ridge and the Hess Rise (the red star in Fig.
1.1), most of the LCDW carried by the deep-circulation current enters the Northeast PacificBasin to the south of approximately 30°N, and the MOC in the Pacific mostly turns over at
mid-latitude region. However, if not, almost the deep water would continue moving to thenorthernmost point at about 55°N in the Northwest Pacific Basin, and then, would enter the
Northeast Pacific Basin along the Aleutian Trench (the blue star in Fig. 1.1), and the MOC in
the Pacific extends to the subpolar region.
Moreover, the author examines variability of the deep-circulation transport in the Pacific.
In many cases, volume transports of the deep-circulation current are estimated withgeostrophic calculation as previously noted. Actually, large variability of volume transport ofthe deep-circulation current is reported from the few mooring results in the Pacific. In theSouthwest Pacific Basin, Whitworth et al. (1999) show its volume transport varying from -6.6
Sv to 44.9 Sv by 22-month moorings (Fig. 1.6) while Tsimplis et al. (1998) estimate its
geostrophic volume transport crossing 32°30′S at 12 Sv by an inverse method. At the Samoan
Passage, Rudnick (1997) shows that the volume transport varies from 1.1 Sv to 10.7 Sv during
17 months (Fig. 1.7) while the mean transport of 6.0 ± 1.5 Sv is similar to geostrophic
transports of 6.0 Sv (Taft et al., 1991) and 7.8 Sv (Roemmich et al., 1996). Kawabe et al.
(2005) estimate the volume transport of the eastern branch passing through the Wake Island
Passage at 3.6 ± 1.3 Sv on an average; it varies from -5.3 to 14.8 Sv (Fig. 1.8). The geostrophic
volume transport of LCDW at the Wake Island Passage is extremely small (nearly zero) in
1999 (Kawabe et al., 2003) when the mooring result shows temporary weakness of current
speed and volume transport (Kawabe et al., 2005). This indicates that snapshots of geostrophic
volume transport from hydrographic sections could not have representativeness and must be
evaluated with direct measurements.
To study these subjects, the author analyzed directly-measured current velocity data andrecent high-quality dissolved-oxygen data. In Chapter 2, the author analyzed current velocity
data from a mooring array with 5 systems from 27°N to 35°N at 165°E from 1991 to 1993 with4

a system at 167°E, 27°N during 1989–1991, and examined the question (3). In Chapter 3, theauthor mapped dissolved oxygen with recent high-quality hydrographic data accumulated since
the WOCE era and examined the questions (1) and (3). In Chapter 4, the author analyzed
current velocity data from a mooring array with 9 systems between the Shatsky Rise and the
Ogasawara Plateau from 2004 to 2005, and examined the questions (2) and (3) as well as thevariability of deep-circulation transport. Finally, the general conclusion is written in Chapter 5.
The contents in Chapters 2 and 3 have been published in Yanagimoto and Kawabe (2007) andthe contents in Chapter 4 have been published in Yanagimoto et al. (2010).

5

Fig. 1.1 Distribution of dissolved oxygen (ml l-1) near the bottom in the Pacific from Mantyla
and Reid (1983). They indicated that the three points shown with blue, red, and green stars are
entrances of the deep water to the Northeast Pacific Basin.

6

Fig. 1.2 Bottom topography and pathways of the deep-circulation current in the western
Pacific. The black arrow shows the deep-circulation current and the following eastern branch,and the gray arrow shows the western branch of the deep-circulation current. Bottom
topography is shown by isobaths of 4000 m. WIP and OP is an abbreviation for the Wake
Island Passage and the Ogasawara Plateau, respectively.

7

Fig. 1.3 Schematic of the deep circulation in the western Pacific by Kawabe et al. (2003).
Black and gray arrows from the Central Pacific Basin show the pathways of the eastern andwestern branch of the deep circulation, respectively.

8

a

b

Fig. 1.4 (a) Distribution of volume transport at a hydrographic section, WHP P10, mostly
along 149°E, and (b) schematic of deep currents at depths of 2500–4500 m based on thesection (Wijffels et al., 1998).

9

Fig. 1.5 Schematic of deep currents based on previous observations and simulations
summarized by Owens and Warren (2001). Numbers indicate references which they adopt (seetheir text).

10Fig. 1.6 Time series of total volume transport at the PCM9 array at 32°30′S and volumetransports of Lower Circumpolar Deep Water (CDW in this figure) and Pacific Deep Water(PDW) (Whitworth et al., 1999).

11

Fig. 1.7 Time series of volume transport deeper than 4000-m depth in the Samoan Passage(Rudnick, 1997).

Fig. 1.8 Time series of volume transport of the eastern branch in the Wake Island Passage(Kawabe et al., 2005). Thin solid and broken line shows the volume transport in the westernand eastern half of the passage, respectively, and thick solid line shows the total volumetransport.
12Chapter 2. Temporal and spatial characteristics ofdeep-circulation current measured at
mid-latitude of 165°E
2.1. Introduction
Lower Circumpolar Deep Water (LCDW) that separates from the Antarctic Circumpolar
Current flows northward along the deep western boundary of the South Pacific, enters theCentral Pacific Basin through the Samoan Passage, and spreads to neighboring basins in the
North Pacific, such as the Northwest Pacific Basin, the Northeast Pacific Basin, and the East
Mariana Basin (Mantyla and Reid, 1983). Figure 2.1 shows topographic features of the western
North Pacific. The pathway of LCDW around the Central Pacific Basin is bifurcated into theeastern and western branches north of the Samoan Passage. The densest, coldest LCDW iscarried by the eastern branch current of the deep circulation and flows northward along theGilbert Ridge at the western periphery of the Central Pacific Basin and through the Wake
Island Passage (Johnson and Toole, 1993; Kawabe et al., 2003). The eastern branch current of
the deep circulation is called the deep-circulation current in the present chapter. The
deep-circulation current through the Wake Island Passage was confirmed by direct current
measurement with a one-year mooring (Kawabe et al., 2005).
In the early 1990s, 165°E meridional hydrographic observations were conducted threetimes as part of the World Ocean Circulation Experiment (WOCE) Hydrographic Programme(WHP): WHP P13C in 1991, P13 in 1992, and P13J in 1993. Kawabe and Taira (1995)
analyzed hydrographic data from WHP P13C and P13J and concluded from geostrophic
velocity and dissolved oxygen distribution that the deep-circulation current proceeds
northward east of 165°E up to 29°N and passes 165°E westward between 32° and 35°N. They
supposed that the deep-circulation current turns cyclonically in the basin between the Shatsky
Rise and the Emperor Seamount Chain. This is based on not only their observations at 165°E
but also the following previous findings: (1) the lowest temperature and nutrients, the highest
salinity and oxygen near the bottom around 165°E in the INDOPAC section at 35°N (Kenyon,1983); (2) the lack of evidence of cold, high-oxygen water north of 36°N in a 165°E
meridional section (Joyce, 1987); and (3) stable northwestward flow at 4000-m depth at 33°N,
165°E measured from 1983 to 1985 with moored current meters (Schmitz, 1987).
Kawabe and Taira (1995) indicated that the deep-circulation current at 165°E changed themeridional location by 1° latitude between WHP P13C and P13J. Figure 2.2 shows dissolved
oxygen distribution on a deep isopycnal surface σ5 = 50.17 kg m-3 observed in three WHP P13

13

sections. The oxygen distribution in WHP P13C and P13J was the same as that in Kawabe andTaira (1995). High-oxygen water was found between 29°N and 35°N, and the location ofoxygen peaks differed among the three sections. This confirms their conclusion of theapproximate position and meridional variation of the deep-circulation current by taking into
consideration the additional data of WHP P13 in 1992, although current variations cannot beclarified with the data collected at the three time points.
In the present chapter, the author reveals the spatial and temporal properties of the
deep-circulation current with analyzing current velocity data obtained with current meters
moored for approximately two years at depths greater than 3800 m at five stations from 27°N
to 35°N, 165°E and at one station at 27°N, 168°E.

2.2. Mooring observations
We conducted direct current measurement for approximately two years using five mooring
systems with 24 current meters at intervals of two degrees of latitude between 27°N and 35°N
at 165°E (Fig. 2.1, Table 2.1). The five moorings were anchored from 20 to 24 August 1991 on
the R/V Hakuho-maru cruise KH-91-5 (WHP P13C) and were recovered from 17 to 22 May
1993 on KH-93-2 (WHP P13J). The moorings were labeled accordingly from M1 at 35°N to
M5 at 27°N. Five current meters were moored near the bottom (30 m from the bottom), one ateach of the five moorings, and three were moored at depths of approximately 4000 m, one eachat M1, M3, and M5. Water depth at every mooring point was approximately 6000 m. Water
depths and mooring depths were based on echo sounding of the ocean floor using a ship-mounted
precision depth recorder (PDR) at the time when the moorings were anchored on KH-91-5. All of

the current meters near the bottom and two current meters at 4000-m depth at M1 and M5
recorded velocity data during the full mooring period. The current meter at 4000 m at M3
failed to function.
Another mooring, CP, was deployed at 27°N, 168°E from 3 November 1989 to 25 August
1991. Mooring CP had three current meters moored at depths of 5820 m (30 m from thebottom), 4820 m, and 3820 m, but the first two current meters failed to function.
We used Aanderaa RCM5 current meters and equivalent instruments manufactured byUnion Engineering Co., Ltd., Japan. Inertial oscillations and diurnal and semi-diurnal tides
were filtered out with a Gaussian filter equivalent to “gaus24” described by Thompson (1983).
We analyzed the values daily at midnight of the coordinate universal time.

2.3. Temporal and spatial variations of near-bottom currents at 165°E
The most prominent feature of the near-bottom current at 165°E is the strong14

northwestward flow with stable current direction at a depth of 6180 m at M2 (33°N), which
continues almost throughout the observation period (Fig. 2.3b). This northwestward flow is themain part of the deep-circulation current (Kawabe and Taira, 1995; Fig. 2.2). The vector mean
velocity is 7.77 cm s-1, 310°T, and is the highest among those of the five moorings (Table 2.1).
This flow has a very sharp peak at 310°T in the current direction histogram and a mean
velocity of 12 cm s-1 for the 310°T sector, which is the highest mean velocity among current
direction sectors (Fig. 2.4). As a result, it has the highest ratio of vector velocity mean to scalarvelocity mean ( | u | | u | ) among the five recorded near-bottom currents (Table 2.1). The ratio
represents stability of current direction; the higher this ratio is, the more stable the current
direction is, and it becomes 1.0 when current does not change direction at all besides
occurrence of no velocity.
The northwestward flow at M2 is strong at WHP P13 (21 cm s-1, 310°T) and P13J (17 cm
s-1, 317°T) at which the northern oxygen peak exists near 33°N, whereas the flow at M2 isweak and southwestward at P13C (6 cm s-1, 239°T), because the deep-circulation current islocated at 33°–35°N (Kawabe and Taira, 1995). Suspensions of the northwestward flow during
August–September 1991, March and June 1992, and December 1992–February 1993,
occurring with intervals of seven or eight months, must be due to a meridional shift of the
deep-circulation current. Therefore, the variations of near-bottom current velocity at M2 aredue to variations of not only velocity amplitude but also position of the deep-circulation
current. Except for the suspensions, the deep-circulation current continues to exist around M2,
keeping the northwestward flow direction, as the major axis of variance is parallel to the meancurrent (Table 2.1).
The relatively high oxygen at 31°N in Fig. 2.2 suggests that the mooring M3 is located in
the deep-circulation current at least temporarily. Near-bottom current at M3 is dominated bynorth-northwestward to northward flows and is almost northward on average, but has
significant velocity in the west-northwestward to northwestward direction despite the low
frequency (Figs. 2.4c and 2.5). The strong northwestward flow during December 1991–January
1992 and April–May 1993 is probably the deep-circulation current that occupies M2 and M3(Fig. 2.3). Thus, the deep-circulation current occasionally expands or shifts southward and islocated around the region 31°–33°N for approximately one month.
Schmitz (1987) showed that a strong northwestward current of 6.2 cm s-1 on average
flowed at 4000-m depth at 33˚N, 165˚E during 1983–1985 (Fig. 2.5). This mean current has
almost the same direction and similar velocity, in comparison with the mean near-bottom
current with a velocity of 7.8 cm s-1 in the present study. The deep-circulation current at
4000-m depth at the Wake Island Passage has a mean velocity of 0.4–0.9 cm s-1 and is much15

weaker than the near-bottom flow with a mean velocity of 1.9–3.5 cm s-1 (Kawabe et al., 2005).
The velocity at 4000-m depth at 33°N, 165°E during 1983–1985 is too high for the
deep-circulation current at 4000-m depth. An anti-clockwise deep eddy would exist south ofthe deep-circulation current as seen in the geostrophic flow distribution at the 165°E section(Kawabe and Taira, 1995), and would overlap the deep-circulation current at 33°N, 165°E.
The near-bottom current at M1 (35°N) has large meso-scale variability, and the current
direction has several peaks in the histogram (Figs. 2.3a and 2.4a). Consequently, the meancurrent speed is very low and not statistically significant (Fig. 2.5, Table 2.1). The frequency ofthe current direction is significant for north-northwestward flow between 330° and 350°T andsouthward flow between 170° and 190°T. The north-northwestward flow during
August–September 1991, during which the deep-circulation current is not located at M2, ispossibly the deep-circulation current. However, the current velocity during this period issimilar to that at depths of 4000 m (Fig. 2.6a) and 2500 m (not shown), having mean velocitiesof 8.47 cm s-1, 329°T near the bottom, 7.99 cm s-1, 333°T at 4000-m depth, and 7.46 cm s-1,
344°T at 2500-m depth. Such a barotropic structure in the deep layer differs from thesignificant upward decrease in strength of the deep-circulation current at the Wake Island
Passage to the extent that the mean velocity at 4000-m depth decreases to one-fourth or
one-fifth of that near the bottom (Kawabe et al, 2005). Moreover, the dissolved oxygen at
WHP P13C is low at M1 (Fig. 2.2). These observations suggest that the deep-circulation
current did not extend to M1, although it was located at 33°–35°N (Kawabe and Taira, 1995).
It is described in Section 2.4 that the deep current at M1 may be due to meso-scale eddies
throughout the mooring period.
Moorings M4 (29°N) and M5 (27°N) are located outside the deep-circulation current. The
near-bottom current at M4 is dominant in the direction of 170°T; it does not have thenorthward component (-90°–90°T) of velocity (Figs. 2.3d and 2.4d) but has a high southward
mean velocity (4.88 cm s-1, 184°T) with high direction stability (Fig. 2.5, Table 2.1). The
near-bottom current at M5 is dominant in the direction of 30°–120°T, with a higheast-northeastward mean velocity (3.13 cm s-1, 75°T) (Figs. 2.3e, 2.4e, and 2.5, Table 2.1).
Mooring M5 is located immediately east of the northern end of a line of deep seamounts
extending northeastward from the Mid-Pacific Seamounts, as shown by the 5500-m isobaths in
Fig. 2.5. The deep flow must be strongly influenced by this seamount, so that the direction of
the near-bottom current at M5 is confined to a relatively small range, and the mean direction(75°T) is similar to the direction of the seamount. The effect of the seamount seems to extend
to the current at 4000-m depth at M5. The direction histogram is similar to that of thenear-bottom current, although the velocity is much lower in all directions (Figs. 2.6b and 2.7b).

16

The current at 4000-m depth at CP (27°N, 168°E) is extremely weak throughout theobservation period and has no directional dominancy (Figs. 2.6c and 2.7c). Its mean velocity is
too low to have any statistical significance (Fig. 2.5, Table 2.1) and is negligible in comparison
with the mean velocities at 4000-m depth at the Wake Island Passage (0.93 and 0.40 cm s-1),
which are statistically significant (Kawabe et al., 2005). This implies that the deep-circulation
current does not pass mooring CP.

2.4. Frequency-domain characteristics of current variations
Power spectral densities of major axis components of velocity variations are shown in Fig.
2.8. The largest power spectral density exists at periods longer than 100 days near the bottom
at M2, having the maximum around 250 days (Fig. 2.8b). The periods of 250 days correspondto the intervals (seven or eight months) of suspensions of northwestward flow at M2, as
discussed in Section 2.3, and the long-term variations may be due to the southward shift of the
deep-circulation current.
Spectral densities over the period of 30 to 80 days at M1 are the largest among meso-scalevariances at all the moorings (Fig. 2.8a). The spectral peaks and magnitudes at periods longer
than 30 days are almost the same among 2500-m, 4000-m, and near-bottom depths at M1.
Coherences between these depths are much higher than the 95% confidence level at almost all
periods longer than 10 days. These suggest that meso-scale eddies having nearly uniform
amplitude and large coherence at depths greater than 2500 m are dominant at M1.
Meso-scale variations over the periods of 50 to 90 days are strikingly dominant near thebottom at M4, being more predominant than those at M1 (Fig. 2.8d). The variations may be
due also to meso-scale eddies at M4. The power spectral density of the near-bottom current at
M5 is depressed between 40 and 100 days (Fig. 2.8e). This contrasts with the meso-scalecharacteristics at M4, although M4 and M5 are close together. The limited meso-scalevariability at M5 may be due to the fact that the variations at M5 are associated with the small
anticyclonic gyre steered by the seamounts, not meso-scale eddies. The power spectrum at
4000-m depth at M5 has a shape similar to that of the near-bottom one, although it is much
weaker. This shows that the variations at M5 are trapped to the bottom, probably by
topographic control of the seamounts.
The power spectrum at 4000-m depth at CP is as weak as that at 4000-m depth at M5 (Fig.
2.8f). The similar energy levels are understandable, since M5 and CP are close to each other.
However, the shapes of their power spectra differ markedly from each other. While the power
spectrum at M5 is depressed between 40 and 100 days, that at CP is nearly uniform at periods
longer than 40 days and decreases sharply from 40 days to 20 days. The power spectra at CP
17

may be composed of the meso-scale variations at M4 and longer-scale variations at M5.

2.5. Conclusion
The present results indicate that the deep-circulation current at 33°N, 165°E is strong andnorthwestward with a mean speed of 7.8 cm s-1 near the bottom, is very stable in the current
direction, and continues for four to six months between interruptions of one or two months‟
duration. Expansion or shift to 31°N, 165°E of the deep-circulation current occurs
intermittently. On the other hand, the deep-circulation current was not recorded, but meso-scale
eddy variations were prominent, at 35°N, 165°E. The deep-circulation current does not reach
35°N, 165°E, although it shifts northward.
The current at 27°N, 165°E has a high east-northeastward mean velocity of 3.13 cm s-1,
75°T with relatively high direction stability near the bottom, while it has a much lower velocity
at 4000-m depth. The shapes of the power spectra of these current velocities strikingly contrast
with those at neighboring moorings in that the power spectral densities are depressed between
40 and 100 days. The bottom-confined current at 27°N, 165°E may be part of the anticyclonicsmall gyre immediately east of the deep seamounts extending northeastward from the
Mid-Pacific Seamounts as mentioned in Chapter 3.

18

Fig. 2.1 Map of the western North Pacific. Regions less than 5000-m depth are shaded. Lines
show isobaths of 4000 m and 6000 m. Stars show positions of five moorings at 165°E and one
mooring at 27°N, 168°E in the present study.

19

Fig. 2.2 Pressure in db (upper panel) and dissolved oxygen concentration in ml l-1 (lower
panel) on the 50.17σ5 isopycnal surface along 165°E, shown by thick solid lines for WHP
P13C in 1991, thin solid lines for P13 in 1992, and broken lines for P13J in 1993. Data points
are marked in the lower panel by pluses (+), crosses (x), and squares (□) for P13C, P13, andP13J, respectively.

20

Fig. 2.3 Stick diagrams of daily velocity near the bottom from August 1991 to May 1993 at
M1 (a), M2 (b), M3 (c), M4 (d), and M5 (e). Three vertical broken lines in the panels show thedates of hydrographic surveys at WHP P13C, P13, and P13J.

21

Fig. 2.4 Direction histogram (dottedcurves) and mean in every 10-degree
directional sector (solid sticks) of daily
velocity near the bottom at M1 (a), M2(b), M3 (c), M4 (d), and M5 (e).
Horizontal and vertical axes show
scales of histogram and speed (cm s-1),
respectively.

22

Fig. 2.5 Mean current velocity vectors (arrows) and variance ellipses. Blue and red arrows
and ellipses show those near the bottom and at 4000-m depth in the present observations,
respectively. Purple ones show those at 4000-m depth for two one-year mooring periods of
1983–1984 and 1984–1985 by Schmitz (1987). Regions shallower than 5000 m are shaded,
and isobaths of 5500 m and 6000 m are shown by thin and thick lines, respectively.

23

Fig. 2.6 Stick diagrams of daily current velocity at a depth of approximately 4000 m at M1 (a),M5 (b), and CP (c). The velocities were recorded between August 1991 and May 1993 in (a)and (b), and between November 1989 and August 1991 in (c).

24Fig. 2.7 Same as Fig. 2.4 but for a
depth of approximately 4000 m at
M1 (a), M5 (b), and CP (c).

25

Fig. 2.8 Power spectral densities of major axis components of current velocity variations near
the bottom (solid line) and at a depth of approximately 4000 m (broken line) at M1 (a), M2 (b),M3 (c), M4 (d), M5 (e), and CP (f).

26

Table 2.1 Statistics of time series of daily current velocities from each instrument at 4000-m
and near-bottom depths shown with suffixes “4” and “B” in record IDs, respectively. Standard
errors of mean current speeds are shown after plus-minus signs. | u | | u | is the ratio of vector
velocity mean to scalar velocity mean, and indicates current direction stability.

RecordMooringID

depth (m)Recordlength(day)

Mean current
Speed(cm/s)Dir.
(°T)

Major axisof variance
(°T)

u
u

M1, 34°59.3′N, 165°09.3′E, Water depth 5858m, from 22 Aug. 1991 to 21 May 1993

M1-4

3930

639

0.30 ± 1.13

208

179

0.06

M1-B

5830

639

0.89 ± 1.38

177

178

0.15

M2, 33°00.1´N, 165°02.3´E, Water depth 6213m, from 24 Aug. 1991 to 19 May 1993

M2-B

6180

635

7.77 ± 1.83

310

135

0.87

M3, 30°59.5´N, 165°00.9´E, Water depth 6051m, from 24 Aug. 1991 to 18 May 1993

M3-B

6020

634

3.57 ± 0.61

4

137

0.72

M4, 28°58.7´N, 165°01.1´E, Water depth 5792m, from 25 Aug. 1991 to 17 May 1993

M4-B

5760

632

4.88 ± 0.37

184

58

0.85

M5, 26°59.9´N, 164°59.6´E, Water depth 5785m, from 26 Aug. 1991 to 16 May 1993

M5-4

4050

630

0.64 ± 0.27

88

81

0.45

M5-B

5750

630

3.13 ± 0.61

75

67

0.62

CP, 26°58.5´N, 167°58.9´E, Water depth 5879m, from 9 Nov. 1989 to 24 Aug. 1993

CP-4

3820

654

0.18 ± 0.23

27

318

75

0.12

Chapter 3. Pathways of the deep-circulation currents in thesouthern part of the Northwest Pacific Basin
3.1. Introduction
The deep-circulation current transporting the oxygen-rich deep water from south, Lower
Circumpolar Deep Water (LCDW), is found to flow north-northwestward at the Wake Island
Passage (Kawabe et al., 2005) and northwestward at 33°N, 165°E (Chapter 2). In order toclarify its pathway, we will examine the deep dissolved-oxygen distribution north of 20°N inthis chapter.
The pathway of the deep water circulation has been studied conventionally with tracing thedistinguishing feature of high dissolved oxygen of LCDW. Mantyla and Reid (1983) drew an
oxygen map near the bottom in the world ocean. They indicate that there is a tongue of highoxygen extending eastward to the Northeast Pacific Basin near Midway Island (the westernend of the Hawaii Ridge) (see Fig. 1.1) and that this would be one of the routes of LCDW
entering the Northeast Pacific Basin. However, this route remains a matter of examining
because their map is based on sparse data which might have been all available at that time.
Kawabe (1993) compiled denser oxygen data in the western North Pacific, which was
accumulated after Mantyla and Reid (1983). His oxygen map on an isopycnal at approximately
4000 dbar shows the western branch of the deep circulation but not the eastern branch. Oxygen
map near 5000 dbar should be drawn to present the deep-circulation current passing throughthe Wake Island Passage (Kawabe and Taira, 1995). At the present day, high-resolution andhigh-quality oxygen data have been accumulated since World Ocean Circulation Experiment(WOCE). In the present study, the eastern branch is examined with oxygen maps below 5000
dbar.

3.2. Data and method
We used WOCE Hydrographic Programme (WHP) data (P03 at 24°N, P02T at 30°N, andP14N at 179°E) and hydrographic data obtained on the R/V Hakuho-maru cruises KH-91-5(WHP P13C), KH-95-1, KH-99-1, KH-03-1, KH-04-4, KH-05-4, and KH-07-1 (Kawabe, 1996,2005a, 2005b, 2006) (Fig. 3.1). Dissolved oxygen concentrations measured with aconductivity-temperature-depth-oxygen profiler (CTDO2) were used except for WHP P03. For
WHP P03, oxygen data at one-decibar intervals were obtained by interpolating water sample
oxygen data with the Akima method. Near-bottom oxygen was calculated by averaging oxygen
data taken 100 m from the bottom, whose depth was estimated with the NOAA National
Geophysical Data Center data set with a 2´ latitude-longitude resolution, usually called
28

ETOPO2.
Water-property data have some gaps among hydrographic sections, as Johnson et al.
(2001) pointed out. Such gaps would include measurement bias and long-term variations.
Since they should be removed to determine synoptic water-property distribution, we adjusted
oxygen and salinity values obtained at WHP P03, P02T, KH-99-1, part of KH-04-4 (12°40′N,
the Wake Island Passage–south of the Shatsky Rise), and KH-07-1 to those of KH-91-5 (WHP
P13C), using values near cross points on an isothermal surface having a potential temperature(θ) of 1.20°C. The adjusted values of WHP P02T were used for adjustment of the observed
values of KH-95-1 and part of KH-04-4 (32°30′N, Shatsky Rise–Izu Ridge). The adjusted
values of WHP P03 were used for the adjustment of WHP P14N. We did not make any
adjustments for KH-03-1, which does not cross any other sections. In accordance with themethod of Johnson et al. (2001), the adjustment values for salinity were added to the original
data, and those for oxygen were multiplied by the original data. The adjustment values for eachsection are listed in Table 3.1. We drew oxygen contours on the near-bottom using 414 data
points and those on 50.17σ5 isopycnal surface using 439 data points, after several data points
were omitted because their oxygen values are extremely large or small among theirsurrounding points in comparison with contour intervals.

3.3. Pathways of the deep-circulation current and dissolved oxygen distributions
The near-bottom oxygen distribution in Fig. 3.2 shows that high-oxygen water in theCentral Pacific Basin mainly spreads northward to the Northwest Pacific Basin through the
Wake Island Passage. It also spreads westward at 13°–18°N but is almost completely blocked
by the seamounts around 163°E. The northward spread forms a north-northeastward
high-oxygen tongue immediately north of the Wake Island Passage, which corresponds to thecurrent direction of 35°T at the Wake Island Passage (Kawabe et al., 2005). The high-oxygen
tongue turns cyclonic and reaches 169.5°–173°E at 24°N, as seen in Roemmich et al. (1991),
168°–169°E at 30˚N, and immediately south of the Shatsky Rise.
Integrating the results of high-oxygen distribution and current properties at CP at 168°E
and M1–M5 at 165°E in Chapter 2 shows the following pathway of the deep-circulation
current in the southern part of the Northwest Pacific Basin. First, it proceeds northward in theCentral Pacific Basin and passes the Wake Island Passage north-northeastward. Then, it passes
24°N at 169.5°–173°N and 27°N around 170°E east of mooring CP. Furthermore, it passes
165°E around 33°N northwestward with meridional variations in the region between north of
30°N and immediately south of 35°N at 165°E. Last, it reaches south of the Shatsky Rise. Thus,
the deep-circulation current forms a cyclonic pathway in the deep basin surrounded by the29

Wake-Necker Ridge, the Emperor Seamount Chain, and the Shatsky Rise.
The oxygen distribution on the 50.17σ5 isopycnal surface (about 5000–6000db) shows thathigh oxygen, >4.00 ml l-1, is confined to the region west of 170°–171°E at 12°–20°N, west of
179°E between the Wake-Necker Ridge and the Hawaiian Ridge, and south of 34°N between
the Shatsky Rise and the Hawaiian Ridge (Fig. 3.3). This suggests that the deep-circulation
current does not bifurcate into an eastward flow in the basin north of the Wake Island Passage,
nor does it flow between the Hess Rise and the Hawaiian Ridge into the Northeast PacificBasin.
The oxygen distribution on the 50.17σ5 isopycnal surface also shows that part of the
deep-circulation current passing 163°–165°E at 18°20´N, as indicated by Kawabe et al. (2003),
proceeds to 20°–21°N, 165°E anticyclonically and probably joins the main part of the
deep-circulation current passing the Wake Island Passage (Fig. 3.3). Moreover, the contour of
3.96 ml l-1 suggests that an anticyclonic small gyre exists with the center at 25°N, 165°E. Thisis consistent with the velocity record at mooring M5 at 165°E shown in Chapter 2. The
east-northeastward bottom-confined current at M5 may form the anticyclonic small gyre
immediately east of the seamounts extending from the Mid-Pacific Seamounts, probably
because of the influence of bottom topography. Thus, the deep-circulation current does notseem to flow westward along the north of the Mid-Pacific Seamounts.
Kawabe (1993), using a map of dissolved oxygen on the 41.52σ3 isopycnal at a depth ofapproximately 4000 m, showed that the western branch current of the deep-circulation flows
northward in the East Mariana Basin. This is not clearly shown in Fig. 3.3, because theisopycnal surface of 50.17σ5 used in Fig. 3.3 is too deep to detect the western branch current.
However, a northwestward extension of the high-oxygen tongue (>4.08 ml l-1) from around
18°N, 155°E to the south of the Ogasawara Plateau and relatively high oxygen east of thisplateau support the pathway of the western branch current inferred by Kawabe et al. (2003).
Moreover, relatively high oxygen around the low-oxygen tongue extending from 12°N, 147°E
to 18°N, 150°E supports the conclusion of Kawabe et al. (2003) that part of the western branchcurrent turns south at the seamounts extending zonally around 20°N, 150°E.

3.4. Conclusion
From the characteristics of current velocity and deep dissolved oxygen distributions, we
specified pathways of the deep-circulation current in the southern part of the Northwest PacificBasin (Fig. 3.4). The deep-circulation current passing mainly through the Wake Island Passage(a thick black arrow in Fig. 3.4) and partially through gaps at 163°–165°E (a thin black arrow
in Fig. 3.4) flows northward at 24°N, 169.5°–173°E and at 30°N, 168°–169°E, proceeds
30northwestward around 33˚N, 165˚E, and goes westward through the south of the Shatsky Rise.
Thus, the deep-circulation current forms a cyclonic pathway in the deep basin surrounded by
the Wake-Necker Ridge, the Emperor Seamount Chain, and the Shatsky Rise. We did not findthat the deep-circulation current proceeded eastward in the region between the Hess Rise andthe Hawaiian Ridge toward the Northeast Pacific Basin. This is confirmed also from thevertical section of dissolved oxygen along WHP P02 at 30°N; the deep dissolved-oxygen
values to the east of 180° is much different from those to the west of 180° (Fig. 3.5). The
wrong suggestion of a deep-water pathway through this region by Mantyla and Reid (1983)
may be due to coarse data points in their map. Relatively oxygen-rich water found to the northof the Hawaiian Ridge would come from the east as part of the deep-circulation current in the
Northeast Pacific Basin (Komaki and Kawabe, 2009; Kato and Kawabe, 2009).
The current axis of the eastern branch seems to pass through the deepest parts in the basinsurrounded by the Wake-Necker Ridge, the Emperor Seamount Chain, and the Shatsky Rise.
The vertical section of dissolved oxygen along WHP P02 at 30°N has a deep maximum core
ranging from the deepest part to the eastern bottom slope (Fig. 3.5). The northwestward
deep-circulation current passes the mooring M2 at 33°N, 165°E, which is in the vicinity of the
deepest part (Fig. 3.6). It is interesting that the eastern branch does not flow along a westernboundary but through the deepest part across isobaths.
The deep dissolved oxygen distributions confirmed the pathway of the western branch current
and its partial turn to the south concluded by Kawabe et al. (2003), and suggested the pathway
shown by a thick green arrow in Fig. 3.4.

31

Fig. 3.1 Location of hydrographic stations used for mapping dissolved oxygen (dots). Pluses
show the moorings of M1–M5 at 165°E and CP at 27°N, 168°E in Chapter 2. Regions
shallower than 5000 m are shaded.

32

Fig. 3.2 Horizontal distribution of dissolved oxygen concentration (ml l -1) near the bottom inthe western subtropical region of the North Pacific. The contour interval is 0.05 ml l-1.
Contours of 4.23 ml l-1 are drawn by broken lines. Hydrographic stations used for this contour
drawing are shown by dots. Thin lines are isobaths of 5000 m.

33

Fig. 3.3 Horizontal distribution of dissolved oxygen (ml l-1) on the 50.17σ5 isopycnal surface
in the western subtropical region of the North Pacific. The contour interval is 0.02 ml l -1.
Contours of 3.95 ml l-1 are drawn by broken lines. Hydrographic stations used for this contour
drawing are shown by dots, and those at which water with 50.17σ5 was not present are shownby crosses. Thin lines are isobaths of 5000 m.

34

Fig. 3.4 Schematic of the deep currents in the western subtropical region of the North Pacific.
The deep-circulation current is shown by black arrows, the western branch current of the deep
circulation is shown by green arrows, and the anticyclonic small gyre around 25°N, 165°E isshown by an orange arrow. Regions less than 5000 m depth and more than 6000 m depth arecolored gray and blue, respectively.

35

Fig. 3.5 A vertical section of dissolved oxygen along 30°N in the North Pacific (Talley, 2007).

36

Fig. 3.6 The zonal (left) and meridional (right) sections of the bottom topography along 33°N
and along 165°E, respectively. The bottom topography is based on ETOPO2. The position ofmooring M2 at 33°N, 165°E is shown by red arrows.

37

Table 3.1

Adjustments of oxygen and salinity to remove gaps among hydrographic sections.

The adjustment values for oxygen are rates to be multiplied by the original data for correction,and those for salinity are differences to be added to the original ones. Sections of KH-04-4
were divided by cruise leg, because the observation areas of the northwestern (leg 1) andsouthern (leg 2) regions are markedly separated as shown in Fig. 3.1.

SectionOxygenSalinity

WHP P02T

1.0215–0.0017

WHP P031.00940.0053

WHP P14N

1.01470.0032

KH-95-1

1.01000.0023

KH-99-1

1.0030–0.0039

KH-04-4 leg 1

1.01870.0012

KH-04-4 leg 2

1.0071–0.0002

KH-05-4

1.02360.0010

KH-07-1

1.00980.000538

Chapter 4. Temporal and spatial characteristics of current
velocities and volume transports ofdeep-circulation currents measured southwest of
the Shatsky Rise
4.1. Introduction
As a part of the global deep circulation, the saline, oxygen-rich, cold Lower Circumpolar
Deep Water (LCDW) is transported northward in the South Pacific by a deep
western-boundary current, entering the Central Pacific Basin mainly through the Samoan
Passage. The LCDW then spreads farther to the Northwest Pacific Basin, the Northeast PacificBasin, and the East Mariana Basin (Mantyla and Reid, 1983). Because of the complex bottom
topography at low latitudes of the western Pacific, the LCDW is transported to the Northwest
Pacific Basin by two branches of a deep-circulation current (Johnson and Toole, 1993; Kawabe
et al., 2003). The deep part of this deep-circulation current forms the eastern branch, flowing
northward along the Gilbert Ridge toward the Wake Island Passage with a volume transport of
8.1 Sv (1 Sv = 106 m3 s–1), whereas the shallow part forms the western branch flowing into the
Melanesian and East Mariana basins with a volume transport of 2–6 Sv (Johnson and Toole,
1993; Kawabe, 1993; Kawabe and Taira, 1995; Kawabe et al., 2003; Kawabe et al., 2006).
The eastern branch partly separates eastward immediately north of 10°N in the Central
Pacific Basin (Kato and Kawabe, 2009) and the remaining part proceeds northward into theNorthwest Pacific Basin through the Wake Island Passage and adjacent gaps in the Mid-PacificSeamounts with a volume transport slightly more than 4 Sv (Kawabe et al, 2003, 2005). The
latter turns cyclonically in the small basin surrounded by the Shatsky Rise, the Emperor
Seamount Chain, and the Wake–Necker Ridge, and then flows westward south of the Shatsky
Rise as shown in Chapter 3 (Fig. 4.1). On the other hand, part of the western branch enters theNorthwest Pacific Basin through gaps in the Magellan Seamounts and the Mid-PacificSeamounts with a volume transport of 2.1 Sv (Kawabe et al., 2003). Then, it detours aroundthe Ogasawara Plateau, which is an eastward bulge of the Izu-Ogasawara Ridge as inferred bydissolved-oxygen distribution in Chapter 3 (Fig. 4.1). Consequently, both branches would passthrough a deep strait (about 950 km width) between the Shatsky Rise and the Ogasawara
Plateau.
To the east of Japan at 38°N, Fujio and Yanagimoto (2005) found westward current on theabyssal plain and northward current on the eastern flank of the Japan Trench; volume transportof the latter current exceeds that of southward current on the western flank. Excess of39

northward transport is also found above the Izu-Ogasawara Trench at 34°N (Fujio et al., 2000).
They pointed out that the existence of the westward current agrees with mooring results ofSchmitz et al. (1987) at 152°E and those of Owens and Warren (2001) southeast of Hokkaido
Island, and presumed that these westward currents would follow the eastern branch and thenorthward excess on the Japan Trench would come from the western branch. Examining thedeep circulation southwest of the Shatsky Rise is also necessary for knowing the downstream
deep circulation in the Northwest Pacific Basin.
We conducted hydrographic and mooring observations along a line located southwest of
the Shatsky Rise. Kawabe et al. (2009) analyzed the hydrographic data in 2004 and 2005, andfound that the lower deep layer where LCDW with potential temperature (θ) lower than 1.2°C
exists is nearly homogeneous. They concluded that geostrophic volume transport carrying
LCDW (θ < 1.2°C) in the lower deep layer is slightly less than 4 Sv and slightly more than 2
Sv for the eastern and western branches of the deep-circulation current, respectively. These
volume transports are consistent with the previous studies. In their results, the location andvolume transport of the branch currents are different between the observations undertaken in2004 and 2005. This suggests temporal variability in the deep-circulation currents. Marked
temporal variability in the deep-circulation currents has been shown with mooring observations
by Whitworth et al. (1999) in the western South Pacific, by Rudnick (1997) at the Samoan
Passage, and by Kawabe et al. (2005) at the Wake Island Passage. In the present chapter, weanalyze the mooring data of current velocity, and document temporal variations in the volumetransport of the deep-circulation currents southwest of the Shatsky Rise.

4.2. Mooring observations
We conducted direct current measurements for approximately 14 months at nine moorings(M1–M9) using 50 current meters along the line from M1 (25°42′N, 149°16′E) to M9(31°13′N, 156°33′E) (Figs. 4.1 and 4.2, Table 4.1). The nine moorings were anchored during
the period 15–23 September 2004 in the R/V Hakuho-maru KH-04-4 cruise and wererecovered during the period 21 November to 1 December 2005 in the KH-05-4 cruise. We used
30 rotor-type current meters manufactured by Union Engineering Co., Ltd., Japan (equivalent
to Aanderaa RCM5), 14 Aanderaa RCM11 Doppler current meters, 5 FSI 3D-ACM acoustic
current meters, and 1 RDI 300-kHz Workhorse Doppler current profiler.
For all nine moorings, we installed current meters near the bottom and at depths ofapproximately 5000, 4500, 4000, and 3500 m. The moorings M1–M3 and M5–M7 wereanchored in a deep basin at water depths of 5700–5900 m, and had additional current meters
installed at depths of approximately 5500 m. The other moorings were located in shallower

40

regions: M4 was anchored on a small seamount at a water depth of 5601 m and had an
additional current meter installed at 2930 m depth, and M8 and M9 were anchored at water
depths of 5112 and 4440 m on the slope of the Shatsky Rise, respectively. The water depths
and depths of current meters listed in Table 4.1 were estimated using depth data converted
from pressure data measured using Aanderaa RCM11 current meters, based on a Sounders andFofonoff‟s method (Fofonoff and Millard, 1983).
Most of the current meters recorded velocity data throughout the full mooring period,
while the current meters at 4890 m depth at M3, 5020 and 4500 m depth at M4, 4930 m depth
at M5, and 4930 m depth at M6 ceased operating during the mooring period, and the current
meter near the bottom at M5 malfunctioned due to a water leak.
Prior to the mooring observations at M1–M9, we deployed two moorings M7p and M8p
during the R/V Hakuho-maru KH-03-1 cruise in May 2003; these sites were located 1.8 km
northeast of M7 and 2.1 km northwest of M8, respectively. Consequently, the water depth at
M8p is nearly 300 m greater than that at M8, although the water depth is similar between M7
and M7p (Table 4.2). The bottom topography around M8p seems to be a small, deep valley
according to the sounding with a ship-mounted precision depth recorder (PDR) of the R/V
Hakuho-maru.
In the following sections, we analyzed daily current velocities subsampled at midnight(coordinated universal time) after filtering out inertial oscillations and diurnal and semi-diurnal
tides using a Gaussian filter equivalent to “gaus24” described by Thompson (1983).

4.3. Velocity field near the ocean bottom
The mean velocity directions below approximately 4000 m depth are within 270°–360°T at
M2, M4, and M6–M9, whereas those are toward 65°T, 163°T, and 244°T at M1, M3, and M5,
respectively (Tables 4.1 and 4.2). Moreover, the direction of the near-bottom current velocity at
M2, M4, and M7–M9 has sharp peaks within the range 270°–360°T in the direction histogram,whereas that at M6 has a dominant peak to the west-southwest, similar to that for M5 (Fig.
4.3).
Because of those velocity characteristics, the vertical section of objectively-mapped meancurrent velocity component normal to the observation line shows that northwestward velocity
components have maximum cores near the bottom at M2, M4, and M8, and the maximum coreat M8 expands toward M7 and M9 (Fig. 4.4). Figure 4.4 was drawn through the following
processes. The obtained current data were mapped to grids of 100 m depth by 0.1° in latitude(about 17 km) for the northeastward observation line with using the optimal interpolation

41

method developed by Bretherton et al. (1976). The optimal interpolation was applied with
influence radii of 0.5° in latitude, 500 m depth, and a noise level of 0.5 cm2 s–2. The sizes of
the influence radii are similar to the spacings of the current meters, and the noise level issimilar to the estimation errors of the mean currents (see Table 4.1).
The pattern of temporal variation in current velocity at M4 is markedly different from thatrecorded at M2 and M7–M9. The northwestward current near the bottom at M4 (red arrows in
Fig. 4.5b) varies significantly between 0 and 20 cm s-1 with a period of 45 days, as shown by
the marked peak in the power spectrum (Fig. 4.6). This current is strongly intensified near thebottom and is highly stable in its direction (Table 4.1). Since M4 was anchored at a water depthof 5601 m on a small seamount, it may be part of a bottom-intensified clockwise eddy around
an isolated seamount located immediately east of M4. The seamount has the top at about 3800
m depth and a horizontal scale of about 20 km (Fig. 4.7). The clockwise eddy is suggested by
the geostrophic flow distribution presented by Kawabe et al. (2009).
Based on the above results, we conclude that the western and eastern branches of the deep
circulation flow northwestward around M2 near the Ogasawara Plateau and at M7–M9 near the
Shatsky Rise.
The eastern branch is almost shown as a current flowing toward 270°–360°T in the stickdiagrams for M7–M9 and M7p–M8p (Fig. 4.5). The current toward 270°–360°T at M8-B
(M8p-40) makes up 77% (80%) of the total daily record, while that at M7-B (M7p-B) andM9-40 accounts for 60% (67%) and 62%, respectively. According to the ratio of the vector
mean to the scalar mean of each velocity ( | u | | u | ), all the current meters at M8 and M9 record
highly stable current directions, being the most stable among the moorings except for thenear-bottom current meter at M4 (Table 4.1). Furthermore, the current meters below 3800 m at
M8 record the highest northwestward mean velocities among the moorings at M7–M9 (Table
4.1). These findings suggest that the eastern branch flows primarily at M8, and secondarily at
M7 and M9.
Current velocity and directional stability of the eastern branch increase toward the bottom
at M7, M7p, and M8 (Tables 4.1 and 4.2); this is similar with the deep-circulation current at
the Wake Island Passage (Kawabe et al., 2005). Exceptionally, current velocity and directional
stability decrease from M8p-40 to M8p-B. This is probably because the current at M8p-50 andM8p-B is affected by the small valley in which M8p was moored. The current at M9-B is also
weaker than M9-40 and M9-35 probably due to disturbance by topography.
The eastern branch was not detected at any mooring during short periods in September
2004 and February–March and May 2005 (Fig. 4.5b), indicating that the eastern branch may
42disappear temporarily, as demonstrated previously in the Wake Island Passage (Kawabe et al.,2005). On the other hand, a current flowing toward 270°–360°T was found at M6 continuously
in November–December 2004 and July–September 2005, indicating that the eastern branch
infrequently expands to M6.
The western branch is detected at M2, being recorded at this mooring from November
2004 to March 2005, from April to June 2005, and from August to September 2005 (Fig. 4.5b).
Currents with northeastward mean velocity at M1 and with southward mean velocity at M3
would not connect with the western branch at M2 because coherences between M1 and M2 andbetween M2 and M3 are not significant at longer periods than 60 days (Fig. 4.8). This suggeststhat the western branch, flowing solely at M2, has a narrower width than the distance between
M1 and M3 of approximately 190 km. The strong current at the peak of current velocity at M2
is not colored in Fig. 4.5b because it is directed toward 0°–10°T; this would undoubtedly be thewestern branch current according to the current direction histogram of Fig. 4.3b. During theremaining periods, the western branch is not apparent in the current-meter records probably
because of its meridional shifts or intermittent disappearance.

4.4. Variability in volume transport of the deep circulation
Daily volume transport of the deep-circulation current was estimated by integrating
velocity components normal to the observation line below the average isotherm of θ = 1.2°C
for the two hydrographic sections obtained during the KH-04-4 and KH-05-4 cruises (theisotherm is shown in Fig. 4.4). We used the objectively-mapped velocities toward 270°–10°T
between M1 and M3 for the volume transport of the western branch and those toward
270°–360°T between M6 and M9 for the eastern branch. The average volume transports of thewestern and eastern branches are 4.1 ± 1.2 Sv and 9.8 ± 1.8 Sv, respectively.
According to the low-passed volume transport composed of 25-day low-passed velocity(Fig. 4.9), the volume transport of the western branch exceeded 5 Sv during the periods
December 2004–March 2005, April–June 2005, and August–September 2005, similar tocurrent velocity at M2, whereas that of the eastern branch exceeded 10 Sv during the periods
October 2004–January 2005, April 2005, and July–November 2005. The volume transport ofthe eastern branch was close to zero during September–October 2004 and February–March and
May 2005, when counter-currents were present at M7-M9 (see Fig. 4.5). Thus, the low-passed
volume transports of the both branches have variations with intervals of 2–4 months. The
temporal variations in volume transport of the western and eastern branches have common
dominant periods of 3 months and 1 month (Fig. 4.10). The 3-month-period variations, which
may correspond to the above 2–4-month-interval one, have a significant coherence of 0.58

43

(95% confidence level is 0.39) between the two branches; volume transport of the westernbranch lags behind that of the eastern branch with a phase of 116° (about 1 month). Although
the volume transport of the western branch has high spectral power at periods shorter than 25days, the large short-term variations may reflect current variations at M2.
The volume transport of the eastern branch increases with increasing its sectional average
current velocity (Figs. 4.9b and 4.9c). The correlation coefficient reaches 0.85. Moreover, its
volume transport also has a high correlation with its current width, yielding a correlation
coefficient of 0.65 (Figs. 4.9b and 4.9d). Daily current width of the eastern branch wasestimated from daily horizontal distribution of volume transport per 1-meter width (Fig. 4.11).
First, positions of the maximum unit-width volume transports were detected between M6 andM9. They form relatively continuous curves as shown by blue lines in Fig. 4.11, except for
short periods in January, February-March, May, and June 2005; the maxima near M5 areseparate from the eastern branch. Then, southwestern and northeastern limits of the easternbranch on the observation line were defined as positions where unit-width volume transport
decreases to one-tenth of the maximum, as shown by red lines in Fig. 4.11. Though thesouthwestern limits are mostly found, the northeastern limits are mostly restricted at M9. The
restriction would not cause much underestimate of current width because the eastern branch is
not expected to expand northeastward further from M9 according to the geostrophic flow by
Kawabe et al. (2009). The average current width of the eastern branch is estimated to be 201 ±70 km.

4.5. Conclusion
Direct velocity measurements with a nine-system mooring array (M1–M9) from 2004 to2005 and additional two moorings (M7p and M8p) from 2003 to 2004 revealed spatial andtemporal properties of the deep-circulation currents southwest of the Shatsky Rise in thewestern North Pacific. The western branch of the deep-circulation current flowing
northwestward (270°–10°T) is detected almost only at M2 (26°15′N) northeast of the
Ogasawara Plateau, having a narrower width than the distance between M1 (25°42′N) and M3(26°48′N) of approximately 190 km. The mean current speed near the bottom at M2 is 3.6 ±1.3 cm s-1. The eastern branch of the deep-circulation current is located at the southwestern
slope of the Shatsky Rise, flowing northwestward mainly at M8 (30˚48′N) on the lower part ofthe slope of the Shatsky Rise with a mean near-bottom speed of 5.3 ± 1.4 cm s-1. It expands toM7 (30˚19´N) on the foot of the rise with a mean near-bottom speed of 2.8 ± 0.7 cm s-1 and M9(31˚13´N) on the middle of the slope of the rise with a mean speed of 2.5 ± 0.7 cm s -1 (nearly
4000-m depth) with a frequency of about 60 % during the mooring period. Furthermore, it
expands to M6 (29˚33´N) twice during the mooring period. The width of the eastern branch is44

201 ± 70 km on average.
Daily volume transports of the western and eastern branch of the deep-circulation current
were estimated with using daily objectively mapped velocities toward 270°–10°T between M1
and M3 and toward 270°–360°T between M6 and M9, respectively. Temporal variations of thevolume transports of the western and eastern branches consist of dominant variations withperiods of 3 months and 1 month, varying between almost zero and significant amount ofvolume transport. The 3-month-period variations are significantly coherent to each other with aphase lag of about 1 month for the western branch. The almost zero volume transport due tothe occupation of counter-currents occurs at intervals of 2–4 months. This finding is similar tothe variations of the current and volume transport at the Wake Island Passage, where thedeep-circulation current weakens and pauses with a counter-current between the peaks of the
four-month variations (Kawabe et al., 2005). In the eastern branch, volume transport is highly
correlated with not only cross-sectional average current velocity but also current width; inother words, volume transport increases with not only current strength but also current width.
The mean volume transport of 4.1 ± 1.2 Sv for the western branch and 9.8 ± 1.8 Sv for theeastern branch is double or more the volume transports of 4 Sv for the eastern branch and 2 Sv
for the western branch, which are estimated by Kawabe et al. (2009) on the basis of theirgeostrophic velocity sections and accumulation of previous results in the upstream regions asreviewed in Chapter 1. This overestimate would be because the current meters are too sparse to
estimate exact volume transport by interpolating the current velocities. The difficulty inproduction of exact current structure may be prominent in wide basins in which the location
and width of current vary over time, like in the strait southwest of the Shatsky Rise. In contrast,
realistic volume transports have been estimated from current-meter data in narrow straits suchas the Samoan Passage (6.0 ± 1.5 Sv; Rudnick, 1997) and the Wake Island Passage (3.6 ± 1.3
Sv; Kawabe et al., 2005). Although the volume transports can be reduced by changing
parameters in the present interpolation method, any realistic velocity sections cannot be
reproduced as is described in Appendix A. On the other hand, particularly in the eastern branch,
the long temporal scale of transport variability such as 3-month period would not differ largely
by interpolation parameters (Appendix A).
At M4 (27°55′N), we found a northwestward current varying with a period of 45 days,
which is probably part of a deep clockwise eddy around an isolated seamount locatedimmediately east of M4. Appendix B shows that a Taylor cap with a clockwise eddy can occur
over a seamount with the similar geographic scale in a steady uniform flow with the typicalcurrent velocity found in the present study. However, it is not clear what causes a meso-scale
periodic oscillation over a seamount. At the southernmost mooring along 152°E by Schmitz et45

al. (1987), which is 28 km northwest of M4, deep current shows a marked periodical variation
with about 45-day period (see Fig. 9f in Schmitz et al., 1982). Imawaki (1985) indicates a
Rossby wave with a period of 44 day and a westward phase speed of 7.5 cm s-1 at fivemoorings around 30°N, 146°40′E, which is located between the Izu-Ogasawara Ridge and thepresent study area. Thus, eddies with 45-day-period variability are frequently seen at theselatitudes in the western North Pacific. The clockwise eddy at M4 might be a trapped one by the
seamount.

46

Fig. 4.1 Location of the nine-system mooring array in the western North Pacific. The „+‟symbols show the positions of the mooring systems (M1–M9). Shading indicates water depths
less than 5000 m and the 4000 m isobaths are shown. Red and blue arrows show the pathwaysof the western and eastern branches of the deep circulation, respectively.

47

Fig. 4.2 Mean velocity distribution over the period 2004–2005. Solid arrows and ellipses
show mean velocities and standard deviations near the bottom, respectively, except for M5 andM9, where those at M5-55 and M9-40 are shown. Dashed arrows show mean velocities atapproximately 4000 m depth. The 5000 and 6000 m isobaths are shown by thin and thick lines,
respectively, and regions shallower than 4000 m are shaded (based on ETOPO2).

48

Fig. 4.3 Direction histograms (black curves) and mean values for 10° intervals (gray lines) of
daily current velocity near the bottom at M1 (a), M2 (b), M3 (c), M4 (d), M5-55 (e), M6 (f),M7 (g), M8 (h), and M9-40 (i). Horizontal and vertical axes show the histogram scale andcurrent speed (cm s–1), respectively.

49

Fig. 4.4 Vertical distribution of cross-sectional components of objectively mapped current
velocity (positive northwestward). Contour interval is 1 cm s–1. The 1.2°C isotherm of potential
temperature is the average of two hydrographic sections obtained during the R/V Hakuho-maru
KH-04-4 and KH-05-4 cruises, for which the station locations are shown by the lower andupper triangles at the top of the panel, respectively. The locations of current meters are shownby diamonds. Bottom topography is based on ETOPO2.

50

a

b

Fig. 4.5 Stick diagrams of 5-day-averaged current velocities near the bottom from 2003 to2004 (a) and from 2004 to 2005 (b). Current velocities at M8p-40, M5-55, and M9-40 are
shown for M8p, M5, and M9, respectively. Orientation is slant so that northwestward flows
perpendicular to the observation line directs upward on paper as shown at the upper rightcorners. Red arrows are velocities toward 270°–360°T.

51

Fig. 4.6 Power spectral densities of the major axis components of current variability near thebottom for M1 to M9 (solid lines) and at 3980 m depth for M4 (dashed line). To eliminate
overlap, each power spectrum is shifted by a 102 (cm s–1)2 (cycle per day)–1 decrement (M3 toM1) or increment (M5 to M9).

52

Fig. 4.7 The narrow, tall seamount with a top at 3800 m depth in the vicinity of the mooringM4. Hatched regions are shallower than 4000 m depth. Contours are shown at every 200 m bybroken lines and at every 1000 m by solid lines.

53

Fig. 4.8 Coherency for major-axis current variations among M1, M2, and M3.

54

Fig. 4.9 Time series of daily and low-passed volume transports of the western branch current(a) and the eastern branch current (b), and time series of cross-sectional average current
velocity (c) and current width of the eastern branch (d).

55

Fig. 4.10 Variance-preserving power spectra of volume transport variability of the westernbranch current (solid line) and the eastern branch current (dashed line).

56

Fig. 4.11 Temporal change of horizontal distribution of volume transport per 1-meter width bynorthwestward velocities in the eastern branch. Contour interval is 50 m3 s-1 and non-zerotransport is shown by hatching. Maximum-transport position is shown by blue lines, and thenortheastern and southwestern limits of the eastern branch are shown by red lines. Distance inkilometers along the observation line from M5 is shown at the top.

57

Table 4.1 Statistics of daily current velocity measured by each moored instrument (September
2004 to December 2005). The suffixes “30,” “35,” “40,” “45,” “50,” “55,” and “B” in record
IDs indicate the record at 3000, 3500, 4000, 4500, 5000, and 5500 m nominal depth and near
the bottom, respectively. Instrument types “UCM”, “RCM11”, “3DACM”, and “ADCP” standfor Union Engineering current meter, Aanderaa RCM11, FSI 3D-ACM, and RDI Workhorse,
respectively. Standard errors of mean current speeds are shown after plus-minus signs. The
value | u | | u | represents the ratio of the vector mean to the scalar mean of velocity (see text).

RecordMooring

InstrumentID

depth (m)type

Mean current

Recordlength(days)Speed(cm s–1)

Dir.
(°T)

Major axis
direction ofvariability
(°T)

u
u

M1, 25°42.1′N, 149°16.3′E, water depth 5760m, from 22 Sep. 2004 to 21 Nov. 2005
M1-35
3290
UCM
424
2.42 ± 1.23
41
47
M1-40
3820
UCM
424
2.84 ± 1.45
58
24
M1-45
4340
UCM
424
3.75 ± 1.47
62
21
M1-50
4860
RCM11
424
3.78 ± 1.67
69
31
M1-55
5380
RCM11
424
4.63 ± 1.78
64
31
M1-B
5710
UCM
424
5.84 ± 1.82
65
37
M2, 26°14.9′N, 149°59.9′E, water depth 5910m, from 21 Sep. 2004 to 23 Nov. 2005

0.55
0.50
0.58
0.50
0.54
0.60

M2-35
3350
UCM
426
2.29 ± 0.72
345
349
M2-40
3870
UCM
426
2.63 ± 0.90
350
347
M2-45
4390
RCM11
426
2.76 ± 1.10
351
348
M2-50
4920
RCM11
426
2.86 ± 1.18
349
348
M2-55
5440
RCM11
426
3.05 ± 1.17
353
352
M2-B
5860
UCM
426
3.56 ± 1.25
347
354
M3, 26°48.0′N, 150°44.0′E, water depth 5780m, from 21 Sep. 2004 to 24 Nov. 2005

0.70
0.71
0.63
0.58
0.58
0.59

M3-35
M3-40
M3-45
M3-50
M3-55
M3-B

0.84
0.79
0.75
0.79
0.78
0.75

3320
3840
4360
4890
5410
5730

UCM
UCM
UCM
3DACM
RCM11
UCM

427
427
427
400
427
427

3.30 ± 0.64
3.18 ± 0.66
3.33 ± 0.77
4.88 ± 0.96
3.54 ± 0.67
3.28 ± 0.65

58

181
171
165
169
170
163

48
54
120
170
177
176

Table 4.1 (continued).
M4, 27°55.0′N, 152°11.5′E, water depth 5600m, from 18 Sep. 2004 to 29 Nov. 2005
M4-30
2930
UCM
429
0.72 ± 0.40
22
351
M4-35
3450
UCM
429
0.84 ± 0.46
8
345
M4-40
3980
UCM
429
1.23 ± 0.56
349
347
M4-45
4500
3DACM
377
1.71 ± 0.73
351
1
M4-50
5020
RCM11
220
2.67 ± 0.76
342
346
M4-B
5550
UCM
429
5.08 ± 0.85347
354
M5, 28°43.7′N, 153°18.1′E, water depth 5830m, from 18 Sep. 2004 to 26 Nov. 2005

0.32
0.35
0.40
0.49
0.81
0.92

M5-35
3360
UCM
431
2.49 ± 0.94
235
259
M5-40
3890
UCM
431
2.97 ± 1.01
241
257
M5-45
4410
UCM
431
2.68 ± 0.94
250
256
M5-50
4930
3DACM
391
1.60 ± 0.74
250
245
M5-55
5460
RCM11
431
3.40 ± 1.02244
240
M6, 29°33.1′N, 154°24.4′E, water depth 5820m, from 17 Sep. 2004 to 27 Nov. 2005

0.60
0.64
0.67
0.37
0.60

M6-35
3360
UCM
433
1.69 ± 1.01
268
240
M6-40
3880
UCM
433
1.70 ± 0.99
259
233
M6-45
4410
UCM
433
1.34 ± 0.64
275
235
M6-50
4930
3DACM
391
2.18 ± 0.84
279
219
M6-55
5450
RCM11
433
2.10 ± 0.86
270
229
M6-B
5770
UCM
433
2.72 ± 0.97284
240
M7, 30°19.0′N, 155°18.4′E, water depth 5705m, from 16 Sep. 2004 to 30 Nov. 2005

0.38
0.40
0.46
0.47
0.48
0.52

M7-35
3380
UCM
437
1.75 ± 1.00
314
321
M7-40
3900
UCM
437
1.86 ± 0.85
306
315
M7-45
4420
RCM11
437
2.10 ± 0.75
303
309
M7-55
5500
RCM11
437
2.34 ± 0.65
311
314
M7-B
5655
UCM
437
2.77 ± 0.72
317
319
M8, 30°47.5′N, 156°00.4′E, water depth 5116m, from 15 Sep. 2004 to 1 Dec. 2005

0.49
0.56
0.59
0.56
0.57

M8-35
3270
UCM
440
2.26 ± 0.53
311
315
M8-40
3800
UCM
440
2.74 ± 0.57
311
310
M8-45
4320
RCM11
440
3.34 ± 0.77
311
309
M8-50
4840
RCM11
440
3.37 ± 0.63
318
335
M8-B
5066
ADCP
440
5.29 ± 1.36
332
332
M9, 31°12.8′N, 156°32.6′E, water depth 4440m, from 15 Sep. 2004 to 1 Dec. 2005

0.83
0.87
0.89
0.85
0.89

M9-35
M9-40
M9-B

0.86
0.85
0.71

3330
3860
4390

UCM
RCM11
UCM

440
440
440

2.55 ± 0.66
2.54 ± 0.71
1.80 ± 0.63

59

283
282
281

286
283
294

Table 4.2 As for Table 1, but for mooring results from May 2003 to September 2004.

RecordMooring

InstrumentID

depth (m)type

Mean current

Recordlength(days)Speed(cm s–1)

Dir.
(°T)

Major axis
direction ofvariability
(°T)

u
u

M7p, 30°19.7′N, 155°19.2′E, water depth 5725m, from 14 May 2003 to 17 Sep. 2004
M7p-302895
UCM
488
0.73 ± 0.32286
300
M7p-403920
UCM
488
1.92 ± 0.59311
314
M7p-50
4945
3DACM
212
2.27 ± 1.06
310
355
M7p-55
5370
RCM11
488
1.84 ± 0.54
311
306
M7p-B
5695
UCM
488
2.38 ± 0.66
314
308
M8p, 30°48.4′N, 155°59.6′E, water depth 5410m, from 14 May 2003 to 18 Sep. 2004

0.56
0.67
0.57
0.62
0.65

M8p-30M8p-40M8p-50
M8p-B

0.78
0.93
0.79
0.60

3205
4230
5255
5380

UCM
UCM
3DACM
RCM11

488
488
215
434

1.18 ± 0.26
3.24 ± 0.55
2.36 ± 0.85
1.09 ± 0.35

60

307
300
349
354

313
294
356
334

Chapter 5. General conclusion
The author clarified the pathway and the volume transport variability of the
deep-circulation current transporting Lower Circumpolar Deep Water (LCDW) in the western
North Pacific with analyzing current velocity data directly measured at two mooring arrays
along 165°E and southwest of the Shatsky Rise. The both arrays are located at key points forthe deep-circulation current in the region between 20°N and 35°N in the Northwest PacificBasin. Five moorings were deployed at 27°–35°N along 165°E from 1991 to 1993 andadditional one mooring at 27°N, 168°E from 1989 to 1991. Nine moorings were deployed
between the Shatsky Rise and the Ogasawara Plateau from 2004 to 2005 and additional two
moorings on the foot of the Shatsky Rise from 2003 to 2004. The author also utilized recent
high-quality and high-resolution dissolved-oxygen data in order to estimate the pathway of the
deep-circulation current between the moorings.
First, the author examined the pathway of the deep-circulation current between 20°N and
35°N in the Northwest Pacific Basin, which has not been confirmed and has been different
among researchers. The points of issue that the author organized the subject on the pathways ofthe deep-circulation current into in Chapter 1 were the following: (1) whether part of theeastern branch flows eastward between the Wake-Necker Ridge and the Hess Rise into theNortheast Pacific Basin or not, (2) whether the western branch proceeds northwestward with adetour east of the Ogasawara Plateau or not, and (3) whether the deep-circulation current flows
westward or eastward between 20°N and 35°N. These questions are answered in the present
thesis as below.
As for the question (1), the author did not find the deep-circulation current flowing
eastward between the Wake-Necker Ridge and the Hess Rise into the Northeast Pacific Basin,
which was supposed as one of the routes of the deep water from south into the Northeast
Pacific Basin based on relatively sparse data of water property (Chapter 3). As for the question(2), the author concluded that the western branch proceeds northwestward with a detour east of
the Ogasawara Plateau (Chapter 3, Chapter 4). As for the question (3), the author concluded
that the deep-circulation current flows not eastward but westward as the eastern branch after
passing the Wake Island Passage (Chapter 2, Chapter 3, Chapter 4).
Thus, after passing north-northeastward through the Wake Island Passage, the easternbranch takes a cyclonic turn, proceeds northwestward around 33°N, 165°E, and takes a detour
south of the Shatsky Rise. Figure 5.1 shows the picture of the deep-circulation current in thewestern North Pacific added the location of the deep current classified by our mooring61

observations at 3 points (marked by blue stars). These findings of pinpoint location could not
be obtained without the mooring observations because the water properties of LCDW of highsalinity and high dissolved oxygen let us know only a rough location with relatively great
width as shown in Chapter 3.
Second, the author examined the temporal variations of volume transports of the westernand eastern branches southwest of the Shatsky Rise (Chapter 4). Time series of their volumetransports have dominant spectral peaks at periods of 3 months and 1 month; the almost zerovolume transport occurs at intervals of 2–4 months. This is the fourth result on the temporalcharacteristics of volume transport of the deep-circulation current in the Pacific, following theresults at the Samoan Passage (Rudnick, 1997), in the Southwest Pacific Basin (Whitworth etal., 1999), and at the Wake Island Passage (Kawabe et al., 2005).
The volume transport variation of the eastern branch flowing northwestward mainly on thelower part of the slope of the Shatsky Rise has rather high correlation with its width variation.
It has an average width of 200 ± 70 km and intermittently expands up to 200 km away
southwestward from the main position. Compared with the deep-circulation current in theSouthwest Pacific Basin with a mean width of 700 km with flow most intense on the flank of
the Tonga-Kermadec Ridge (Whitworth et al., 1999), the eastern branch in the Northwest
Pacific Basin has only about one-third width. The western branch is also so narrow that we
find it almost exclusively at one mooring northeast of the Ogasawara Plateau.
These findings on the temporal and spatial structures of the deep-circulation current
between 20°N and 35°N in the western North Pacific do not only elucidate the deep circulation
pathway that has been different for researchers so far, but also suggest the northward extensionof the meridional overturning circulation (MOC) in the Pacific Ocean. The two branches of the
deep-circulation current would converge to the east of Japan (Fujio and Yanagimoto, 2005),
and the confluence flows northward east of the Japan Trench at 40°N (Kojiro Ando, personalcommunication). The deep circulation would continue flowing northward along the Japan and
Kuril trenches to the northernmost point of the Northwest Pacific Basin at about 55°N. Thus,
almost all LCDW entering the Northwest Pacific Basin flows into the Northeast Pacific Basinthrough the Aleutian Trench region. This leads to the conclusion that the MOC in the Pacific
Ocean extends to the Aleutian Trench at 55°N, namely, further north than the northern limit(20°–40°N) of the MOC represented from geostrophic velocities at several transpacifichydrographic sections with the inverse technique by Macdonald et al. (2009) and so on.
The present thesis demonstrates two subjects which should be clarified in the next step ofdeep circulation studies. The first subject is the cause of the 3-month-period variation in62

volume transport of the deep-circulation current southwest of the Shatsky Rise. We compare
volume transport variations along the pathway of the deep-circulation current shown in Fig.
5.2d. The 3-month-period variation of volume transport of the eastern branch is similar to thatat the Wake Island Passage (Kawabe et al., 2005) (Fig. 5.2a), while it is not similar to those at
the Samoan Passage (Rudnick, 1997) (Fig. 5.2b) and in the Southwest Pacific Basin(Whitworth et al., 1999) (Fig. 5.2c). Rudnick (1997) speculated that the 1-month-periodtransport variation at the Samoan Passage would be due to a resonance in the Samoa Basinsouth of the passage. Kawabe et al. (2005) speculated a relation of the 3–4-month-periodvariation at the Wake Island Passage to the dominant 90-day-period variation in sea level at the
Wake Island, which is due to westward propagation of the disturbances generated at the Big
Island of Hawaii (Mitchum, 1995). The 3–4-month-period transport variation of the easternbranch may be propagated downstream and found southwest of the Shatsky Rise.
The second subject is the dynamical mechanism of pathway of the eastern branch entering
the Northwest Pacific Basin through the Wake Island Passage. The eastern branch current
passes the deepest part of the small basin surrounded by the Wake-Necker Ridge, the Emperor
Seamount Chain, and the Shatsky Rise, turns anticlockwise in this basin, and reaches
immediately south of the Shatsky Rise. This shows that the eastern branch in the Northwest
Pacific Basin is not the western boundary current as it does not flow along the Mid-PacificSeamounts. The other parts of the deep-circulation current flow northward along the westernboundary of the Pacific Ocean. In a 1-degree-grid numerical model, Ishizaki (1994) showedthat the eastern branch after passing the Wake Island Passage flows partly along the
Mid-Pacific Seamounts and partly northward along the Shatsky Rise with turning clockwise in
the small basin. If the bottom topography is expressed more realistically in a finer-grid modelwith 11-km grids, the eastern branch flows not along the Mid-Pacific Seamounts but along the
f/h contours, and is modified more realistically by taking upwelling at the current into
consideration (Shinzou Fujio, personal communication).
Dynamics of these phenomena should be examined using ocean models. Before that, they
may have to be described more exactly by further observations.

63Fig. 5.1 The pathways of the deep-circulation current in the western North Pacific. The blue
arrow shows the eastern branch, and the orange arrow shows the western branch. The positionsof these branches in the Northwest Pacific Basin are identified by the moorings marked by bluestars in the present thesis. Bottom topography is shown by isobaths of 4000 m and shades forshallower than 5000 m.

64

a

b

c

d

a

b

c

Fig. 5.2 Power spectrum of volume transport of the deep-circulation current (a) at the Wake
Island Passage (thick line) (Kawabe et al., 2005), (b) at the Samoan Passage (Rudnick, 1997),
and (c) in the Southwest Pacific Basin (denoted by CDW) (Whitworth et al., 1999). The map(d) shows the positions of (a)–(c) by red stars and the positions of the moorings in the present
study by blue stars. Black arrows are the deep circulation current and the following easternbranch and gray arrows are the western branch in the map (d).

65Appendix A. Dependency of volume transport estimationfrom current meter data on interpolation
methodIn order to estimate daily volume transports southwest of the Shatsky Rise fromcross-sectional components of directly-measured current speeds at rather sparse mooring
systems, the author integrated daily grid-point components interpolated to fine grids with avertical interval of 100 m and a latitudinal interval of 0.1° on the mooring-array section(Chapter 4). The author used a traditional optimal interpolation method by Bretherton et al.
(1976), which needs a correlation matrix for state estimation constructed from a priori
parameters: a noise level and horizontal and vertical influence radii. This optimal interpolation
method does not only interpolate between data points but also smoothes observed data within
the noise level. Dependency of volume transports of the eastern and western branches on theseparameters is examined as in Table A.1. Within this range of parameters, the volume transportof the eastern branch changes from 11.8 ± 2.2 Sv in Case c1 to 5.0 ± 0.8 Sv in Case a3, andthat of the western branch changes from 5.0 ± 1.4 Sv in Case b1 to 2.0 ± 0.5 Sv in Case c3.
Error ranges of volume transports in Table A.1 are merely standard errors in estimating meanvalues from time series of transports. Differences among the cases are generally larger thanthese standard errors although the range of parameters can be selected arbitrarily. The larger
noise level is given, the smaller interpolated speeds are located at grid points, and then, thesmaller volume transports are obtained (Cases 1 to 3). Dependency on influence radii (Cases ato c) is rather complex for the western branch while the transport of the eastern branchdecreases with an increase in influence radii.
Case b2 is the parameter combination leading volume transports of 9.8 ± 1.8 Sv for theeastern branch and 4.1 ± 1.2 Sv for the western branch in Chapter 4. In this case, the noiselevel is comparable to the standard errors of the mean velocities and influence radii are similar
to the spacing of the mooring instruments. Furthermore, the mean velocity section in Case b2
seems more realistic than those in Cases 1 and 3 and Cases a and c (Fig. A.1). The velocity
section in Case a3 is much different from our familiar one; there are many velocity cores. The
velocity section in Case c3 is too strongly smoothed; it has a too wide negative core near 27°N.
Although we have no objective criterion to discriminate realistic velocity sections, this optimal
interpolation method is unlikely to lead realistic volume transport, which is similar to thetransports in Cases a3 and c3, within parameters producing realistic velocity sections. A simple
linear interpolation also leads large volume transport of 11.1 ± 2.1 Sv for the eastern branch
and 4.3 ± 1.3 Sv for the western branch.

66

Because the gridded current speeds are small in Cases a3 and c3, zero volume transports
occur more often than in Case b2 (Fig. A.2). However, rough shape of the time-series diagramof volume transport of the eastern branch in Cases a3 and c3 does not differ much from that inCase b2 (Fig. A.2a). On the other hand, that of the western branch largely varies withinterpolation parameters (Fig. A.2b).

67

Table A.1 Volume transport depending on the parameters of objective interpolation method ofBretherton et al. (1976). Upper and lower value in each column is volume transport by Sv (1
Sv = 106 m3 s-1) of the eastern branch and the western branch, respectively. Rx and Rz indicateshorizontal influence radius (latitudinal degree) and vertical influence radius (meter),respectively.
Noise level

Cases 1
2

Cases 22

Cases 3-2

4.0 cm2 s-2Case a1
7.7 ± 1.3 Sv
4.2 ± 1.0 Sv

Case a2
6.9 ± 1.1 Sv
3.7 ± 0.8 Sv

Case a3
5.0 ± 0.8 Sv
3.2 ± 0.5 Sv

Case b1
10.6 ± 2.0 Sv
5.0 ± 1.4 Sv

Case b2
9.8 ± 1.8 Sv
4.1 ± 1.2 Sv

Case b3
6.8 ± 1.2 Sv
2.8 ± 0.6 Sv

Case c1
11.8 ± 2.2 Sv
4.6 ± 1.4 Sv

Case c2
11.2 ± 2.1 Sv
2.9 ± 0.9 Sv

Case c3
8.8 ± 1.7 Sv
2.0 ± 0.5 Sv

0.1 cm s-2

0.5 cm s

Influence radii

Cases aRx = 0.25°, Rz = 250m

Cases bRx = 0.5°, Rz = 500m

Cases cRx = 1.0°, Rz = 1000m

68aDepth (m)Latitude (°N)

b
Depth (m)Latitude (°N)

c

Depth (m)Latitude (°N)

Fig. A.1 Mean normal velocity sections southwest of the Shatsky Rise in Case b2 (a), Case a3(b), and Case c3 (c). Panel a is the same as Fig. 4.4 in Chapter 4. Contour interval in Panel a, b,
and c is 0.5 cm s-1, 0.1 cm s-1, and 0.25 cm s-1, respectively.

69

a

b

Fig. A.2 Time series of volume transports smoothed with 25-day low-pass filter for the easternbranch (a) and the western branch (b). Red, green, and blue lines in each panel are the timeseries in Cases b2, a3, and c3, respectively. Red lines for Case b2, based on the parameters
adopted in Chapter 4, are the same as thick lines in panels a and b of Fig. 4.9.

70

Appendix B. Anticyclonic eddy over an isolated seamount
Geostrophic flow in ideal homogeneous fluid follows the isobaths with avoiding an
isolated topographic obstacle, over which water is trapped. This trapped-water region is called
Taylor column. On the other hand, stratification permits geostrophic flow to pass over an
obstacle, and the trapped-water region becomes confined to the vicinity of the obstacle with ashape like cap or cone. This is called Taylor cap or Taylor cone (Hogg, 1973). Various aspectsof Taylor caps have attracted many researchers for a long time; particularly from the 1970s to
the 1990s, many observational studies were made for documents on Taylor-cap-like effects andmany theoretical studies with numerical and laboratory experiments were made forunderstanding the dynamics of Taylor caps. Roden (1987) reviews theoretical studies mainly
based on quasi-geostrophic approximation and observational studies in the 1970s and the1980s. Numerical modeling has evolved into primitive equation model; Chapman and
Haidvogel (1992) present the conditions of formation of Taylor caps over seamounts in steady,rotating, stratified flows.
In a steady, rotating, nearly inviscid, stratified flow, the following parameters generally
controls aspects of the flow over a seamount:
Rossby number: Ro = U/fL,
Burger number: S = NH/fL,
Beta parameter: b = βL2/U,
Fractional seamount height: δ = hm/H,
where U: the speed of the spatially uniform inflow, L: the horizontal scale of the seamount, hm:
the seamount height, H: the depth of the flat bottom around the seamount, N: the buoyancy
frequency, f: the Coriolis parameter, and β: the beta parameter.
The overview of transient response of the flow over a seamount is expressed as thefollowing (McCartney, 1975; Verron and LeProvost, 1985; Chapman and Haidvogel, 1992). A
spatially uniform inflow impinging on the seamount causes upwelling on its upstream side anddownwelling on its downstream side. Stratification parameterized by burger number controls
the upwelling and downwelling. The upwelling and downwelling form anticyclonic andcyclonic eddies by shrinking and stretching of water column, respectively. In the Northern
Hemisphere, these eddies co-rotate anticyclonically around the seamount with advected by the
inflow. Then, the anticyclonic eddy moves to the top of the seamount and is trapped over the
seamount, and the cyclonic eddy is carried downstream (Verron and LeProvost, 1985;
Chapman and Haidvogel, 1992). If the inflow is strong enough (Rossby number is large), the71

anticyclonic eddy is also carried downstream and a Taylor cap is not formed (Chapman and
Haidvogel, 1992). If the seamount has an enough large horizontal scale, we must consider thebeta effect (Beta parameter is large). In this case, the cyclonic eddy carried downstream by theeastward inflow remains by Rossby wave and develops a complex flow pattern to the east of
the seamount. On the other hand, the cyclonic eddy carried downstream by the westward
inflow does not develop. Thus, the beta effect causes an asymmetric response to the directionof zonal inflow (McCartney, 1975). If the fractional seamount height is large enough (alsostratification is present), the upstream upwelling moves lower dense water to higher layer. This
strengthens local density stratification and causes bottom intensification of Taylor cap(Chapman and Haidvogel, 1992).
The seamount near the mooring M4 southwest of the Shatsky Rise in this study (Chapter4), which is at 28°N (f = 7 × 10-5 s-1, β = 2 × 10-3 m-1 s-1), has a horizontal scale (L) of 20 km
and a height (hm) of 2200 m (the ambient water depth is ~6000 m and the top is at the depth of~3800 m) (see Fig. 4.7). Our hydrographic observations southwest of the Shatsky Rise indicate
deep density stratification with the buoyancy frequency (N) of 3 × 10-4 s-1. Burger number (S)is about 1.3, which means the deep stratification in this study area is comparable with theeffect of the earth rotation. If the inflow has a speed (U) of 5 cm s-1, which is similar to (orrather higher than) the typical current speed in the deep layer of this study area, Beta parameter
becomes 0.16 and we can ignore the beta effect. Therefore, we examine the condition offormation of a Taylor cap over the seamount on an f-plane with using the diagram of Chapman
and Haidvogel (1992), which shows regimes of Taylor-cap formation by Rossby number (Ro)and fractional seamount height (δ) (Fig. B.1). Their setting of Burger number is S=1, which isalmost the same as that in this study area. In the regime of Taylor-cap formation, waterparticles remain over the seamount with rotating anticyclonically as in Fig. B.2a, while they
are swept downstream in the regime of “TT” with slightly larger Rossby number (Fig. B.2b).
In the case of this seamount,
Rossby number: Ro = U/fL ~ 0.04,
Fractional seamount height: δ = hm/H ~ 0.37.
This combination of parameters is within the regime of Taylor-cap formation.

72

Fig. B.1 Regimes for steady, stratified flow (Burger number is 1) over a seamount estimatedfrom the numerical modeling (divided by solid lines) (Chapman and Haidvogel, 1992). “TT”
means the regime of temporarily-trapped eddy before being swept downstream. Regimes
divided by a broken line are predicted from quasi-geostrophic (QG) theory.

73

Fig. B.2 Plan views of trajectories of two water particles departed from near-bottom in steady,stratified flow (S = 1) over a seamount with a moderate height (δ = 0.6) (Chapman and
Haidvogel, 1992) with inflow speed of Ro = 0.1 (a) and Ro = 0.12 (b). Inflow is from left toright. Particle positions at every 5 days are shown by symbols. Larger symbols denote the startpoints.

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