Evidence of Vertical Coupling between the Kuroshio ... - po.gso.uri.edu

Journal of Marine Research, 70, 719–747, 2012

Evidence of Vertical Coupling between the Kuroshio Extension and Topographically Controlled Deep Eddies by Andrew D. Greene1,2,3 , D. Randolph Watts4 , Georgi G. Sutyrin4 and Hideharu Sasaki5 ABSTRACT Strong energy in the 30–60 day band was observed using 39 deep pressure and current records from the Kuroshio Extension System Study (KESS). Energy in this band accounted for 25–50% of the total deep-pressure variance and was strongest under the Kuroshio Extension jet axis. Often, deep-pressure anomalies propagated into the region from the north-northeast and locally intensified as they passed under and interacted with the Kuroshio Extension. The topographically controlled deeppressure anomalies translate nearly along lines of constant f/H . Statistically significant coherence between 30–60 day upper- and deep-ocean streamfunction anomalies demonstrated that there was strong vertical coupling in that time band. Twenty-five percent of the total upper-ocean streamfunction variance was contained within the 30–60 day band near the Kuroshio Extension. Joint CEOFs of the upper- and deep-ocean streamfunctions revealed that near the axis of the Kuroshio Extension the phases were laterally offset alongstream, with the deep ocean leading the upper ocean. This arrangement is attributed to producing joint development of upper-ocean meanders and deep-pressure anomalies. A numerical process model simulated the interaction of barotropic TRWs with an eastward-flowing baroclinic jet. When the TRWs, used as a surrogate for topographically steered deep-pressure anomalies, passed under the jet, they intensified and upper-ocean meanders steepened, much like the observed interactions. The model illustrates how the interaction between TRWs and an eastward-flowing jet, at its simplest level, can reproduce many of the major traits of our observations. The Ocean General Circulation Model for the Earth Simulator also showed similar processes in the 30–60 day band in the KESS region. The strongest variance in the deep fields occurred under the Kuroshio Extension. Upper and deep low- and high-pressure anomalies propagated south southwestward across the Kuroshio Extension, with model phase speeds and wavelengths matching the KESS observations.

1. Introduction In 2004 the multi-institutional collaborative project titled the Kuroshio Extension System study, KESS, was launched. KESS aimed to understand the processes that link upper and 1. Naval Undersea Warfare Center Division Newport, Newport, Rhode Island, U.S.A. 2. Current address: Naval Undersea Warfare Center Division Newport, 1176 Howell St., Newport, Rhode Island, 02841, U.S.A. 3. Corresponding author e-mail: [email protected] 4. Graduate School of Oceanography, University of Rhode Island, Narragansett, Rhode Island, U.S.A. 5. Japan Agency for Marine-Earth Science and Technology, Yokohama, Japan. 719

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deep circulation in the Kuroshio Extension. One working hypothesis, motivated by previous Gulf Stream studies, was that meanders of the upper baroclinic jet in the Kuroshio Extension produce coupled deep-pressure anomalies via baroclinic instability (Shay et al., 1995; Savidge and Bane, 1999). In addition to local development of meanders and deep-pressure anomalies, this paper reports evidence that remotely generated nearly depth-independent eddies propagate across the KESS region and repeatedly develop into intensified deeppressure anomalies while a joint steepening of upper-ocean meanders occurred. After initiation by these incoming eddies, the subsequent intensification process of deep-pressure anomalies and meanders is consistent with a baroclinic instability growth process. Early Gulf Stream studies that highlighted the time-varying nature of the Gulf Stream path characterized meander wavelengths and periods, constructed dispersion diagrams, and calculated temporal and spatial growth rates of propagating meanders (Fuglister and Worthington, 1951; Hansen, 1970; Watts and Johns, 1982; Tracey and Watts, 1986; Lee and Cornillon, 1996). The Synoptic Ocean Prediction Experiment (SYNOP) demonstrated that developing Gulf Stream meanders coupled dynamically to coherent deep-pressure anomalies and involved the full water column (Shay et al., 1995; Watts et al., 1995). In steep meanders, Shay et al. (1995) showed that over time-scales of 30–60 days, Gulf Stream meander troughs (crests) developed and intensified jointly with deep cyclones (anticyclones). Case studies of steep-meander events revealed that the spin-up of deep cyclones and positive relative vorticity generation could be accounted for by the local meander-induced stretching of the lower water column (Savidge and Bane, 1999; Howden, 2000). These observations confirmed the theoretical notion that baroclinic instability is the primary mechanism for upper-ocean meander growth and the generation of deep-pressure anomalies (Pedlosky, 1987). Theoretical studies, baroclinic and barotropic, on jet meandering have taken four seemingly distinct approaches discussed by Flierl (1999). Thin jet models attempt to build a path equation which will predict the evolution of the jet axis from its initial position. Contour dynamical models provide more complicated path equations derived from piecewise constant potential vorticity profiles and have been used as tools for understanding the processes of eddy formation and meander growth. Instability theory focuses on understanding the growth and propagation of waves on the jet. Finally, numerical process models allow simulations of the jet without the severe approximations inherent in the previous methods, but require detailed information on initial conditions and boundary conditions, information which is often unavailable. Topography can also influence the stability properties of a jet. A slope upward to the north across the jet stabilizes the jet (e.g., Sutyrin et al., 2001) while along-jet variations of meridional slope provide a mechanism for local baroclinic instability (Bracco and Pedlosky, 2003). Closer to Cape Hatteras, upstream of the region of steepest Gulf Stream meanders, the deep variability has the character of topographic Rossby waves (TRWs) with periods ranging from eight–64 days, strongest near 40 days (Hogg, 1981; Johns and Watts, 1986; Pickart and Watts, 1990; Hamilton et al., 1996). Rizzoli et al. (1995) suggested that TRWs could

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be forced by eastward-propagating meanders if topographic beta (β) is sufficiently large compared to planetary β as to rotate the Rossby wave dispersion relation and permit an eastward component of phase speed. The resonance condition for coupling is satisfied if Gulf Stream meanders have frequency and eastward wavenumber (ω, k) matching that of an allowed TRW. The generalized resonance condition for non-zonally propagating meanders is (ωmr , kmr ) = (ωtr , ktr ), where kmr and ktr are the rotated components of the meander and TRW wavenumber, respectively, along the local path of the Gulf Stream. Pickart (1995) found that along the Gulf Stream northeast of Cape Hatteras, meridional topographic slopes do in fact dominate planetary β, and eastward-propagating meanders with ≈ 40 day period could couple to and account for the observed TRWs in the SYNOP Inlet array. Farrel and Ioannou (1996) summarize a suite of atmospheric studies in which preexisting finite amplitude perturbations initiate a “transient growth” process, capable of producing strong meanders of the jet stream and intense lower-atmosphere cyclogenesis/anticyclogenesis within time intervals of the same order as the wave period. They point out that the observed spatial and temporal scales of instabilities are difficult to reconcile with those predicted from classic baroclinic instability theory (e.g., Eady, 1949). For instance, most cyclogenesis events in the atmosphere take place on 12–48 hour time scales and therefore there is a need to study the transient growth of finite perturbations, rather than exponential growth from fastest growing but infinitesimal perturbations, allowing time to approach infinity. Traditional studies of baroclinic instability are inappropriate because they seek to find those fastest growing linear normal modes, whereas in many cases finite amplitude perturbations with suboptimal growth rates are found to have produced the observed meanders and lower-atmosphere pressure-anomaly systems. Analyses of the KESS data set presented here reveals the presence of strong energy in the 30–60 day time band in both the upper and deep ocean. Deep-pressure anomalies propagate south-southwestward (SSW) and interact with the upper Kuroshio Extension jet. Bishop et al. (2012) showed that KESS deep-pressure anomalies are associated with weakly bottomintensified (8.3 km vertical scale) topographically-controlled eddies. Here deep-pressure anomalies are conceptualized simply as weakly depth-dependent or “barotropic.” The passage of a deep anticyclone (cyclone) under the Kuroshio Extension results in intensification of a Kuroshio crest (trough) while also providing additional local spin-up of the deep anticyclone (cyclone). A conceptual model and an idealized numerical simulation is then used to elucidate the interaction process between finite amplitude southward-propagating deeppressure anomalies and an upper-ocean eastward flowing jet. Lastly, data from the Ocean General Circulation Model for the Earth Simulator are analyzed, revealing that upper-ocean meanders steepen and deep-pressure anomalies intensify as they propagate underneath the Kuroshio Extension just as in the KESS observations. This paper is organized as follows. Section 2 describes the observational data set. Section 3 explains the methodologies used within the manuscript. Section 4 shows the results of various time and spatial data analysis techniques. Section 5 provides a discussion of the observations as well as results from an idealized numerical model of the interaction between

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Figure 1. KESS observational array. CPIES are denoted by solid black diamonds. Open circles identify the tall current meter moorings locations. Superimposed is the mean altimetric sea surface height (cm) between 1992–2006 using AVISO Rio05 as the mean field. Topography between 0–4,000 m is shaded in light grey colors. Sea surface height standard deviation >31 and 39 cm is color shaded in yellow and salmon, respectively.

barotropic topographic Rossby waves and an eastward-flowing upper-ocean jet. Section 5 also examines 30–60 day deep- and upper-ocean anomalies from a circulation model, the Ocean General Circulation Model for the Earth Simulator (OFES). Section 6 offers conclusions and suggests future work. 2. Observations and numerical models Forty-six current and pressure recording inverted echo sounders (CPIES) were deployed in the Kuroshio Extension region, 32–37◦ N and 143–149◦ E, during May 2004–July 2006 (Fig. 1). They were moored in depths ranging from 5,300–6,400 m. The meridional span of the CPIES array ensured that the meandering Kuroshio Extension front was fully encompassed within the array, while the zonal extent aimed to capture the quasi-stationary meanders at 144◦ and 147◦ E. The CPIES array was also embedded in the regional maximum of sea-surface height variability, represented by the yellow and salmon shading in Figure 1. Instruments were spaced on average 88 km zonally along lines separated 88 km meridionally, diagonally offset. Koblinsky et al. (1984) showed that the isotropic spatial correlation length scale was 90–100 km for temperature at 300 m depth in the region, therefore our array spacing was well suited for studying mesoscale processes. For a comprehensive discussion

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on the latest CPIES processing techniques, assessment of the KESS array’s instrument performance, and how CPIES data products (e.g., temperature and velocity) are computed and compared with in-situ measurements, the reader is referred to Donohue et al. (2010) and Kennelly et al. (2008). The inverted echo sounder (IES) component measures the vertical acoustic travel time (τ) round trip from the seafloor to the surface and back. Variations in τ arise mainly due to changes in thermocline depth. Each IES was equipped with a pressure sensor (PIES) to measure deep pressure variations (p). Thirty-seven PIESs also had an AanderaaTM Doppler current sensor tethered 50 m above the bottom (CPIES). All the measured variables, τ, p, u, and v, were processed to an hourly output interval, where u and v are the zonal and meridional velocities measured by the current sensor. In addition, there were seven deep moorings with RCM-11 current meters, sampling at 5,000 m, deployed across the Kuroshio Extension. All the data from the KESS array were low-pass filtered using a fourth-order Butterworth filter with a three-day cutoff period. This filtering removes the major tidal constituents, inertial oscillations, and the scatter in τ produced by surface waves, while retaining the low-frequency mesoscale variability. Transients in the filtered time series were reduced by dropping the first and last day (one-third of the filtering period) of each time series (Kennelly et al., 2008). Of the 46 CPIES, 39 operated continuously between June 1, 2004 and June 8, 2005. This time interval chosen, for time series analysis, was guided by a balance between maximizing both the time series length and the number of operating instruments. The idealized numerical process model used is governed by two-layer “intermediate” geostrophic equations of motion, guided by Sutyrin et al. (2003). Intermediate refers to a simplified set of equations that are between full primitive and quasi-geostrophic, in which the flow is geostrophic and depth variations within each layer are not constrained to be of order Rossby number times the water column depth. The model domain is 5,120 km × 2,560 km with 10 km grid resolution. The intermediate equations and numerical-solving algorithm are outlined in Sutyrin et al. (2003). The Ocean General Circulation Model for the Earth Simulator (OFES), operated by JAMSTEC, was also used for comparison with KESS observations. OFES is based on the MOM3 platform with 0.1◦ horizontal resolution and 54 vertical levels. Bottom cell construction takes advantage of the partial cell method (Pacanowski and Gnandesikan, 1998) which allows for variable bottom thickness of the deepest vertical level, producing realistic bottom topography effects, e.g., topographic Rossby waves. After a 50-year climatological spin-up (Masumoto et al., 2004), the model is forced by daily-mean NCEP/NCAR reanalysis wind stress and surface tracer fluxes for the years 1950–2006 (Sasaki et al., 2008).

3. Methods Measured travel times were calibrated with coincident CTD casts and leveled to the 1,400 dbar level. The “leveling” process eliminates inter-site path-length effects due to

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Figure 2. Relationship between travel time τ and geopotential anomaly φ illustrated from historical hydrography. τ is integrated from CTD data from the surface to 1,400 dbar. The purpose of this is to show the strong relationship between τ and φ. The grey curve is a cubic spline and is used to covert an IES τ measurement to φ.

varying instrument depths so that τ anomalies can be combined among sets of instruments and interpreted dynamically. Calibration is equivalent to knowing the instrument pressure depth exactly. The calibration and leveling procedure is discussed in Kennelly et al. (2008) and Donohue et al. (2010). Fifty-four Argo float profiles (www.usgodae.org) within 10 km of an IES provided independent estimates of τ and when compared to IES τ yielded an rms difference of 0.95 ms, consistent with measurement error predictions (Donohue et al., 2010). Watts and Rossby (1977) demonstrated that τ could be used to estimate geopotential height φ. We used historical hydrography from the Kuroshio Extension region to establish an empirical relationship between τ1400 and φ5300 (Fig. 2). The rms error in using τ1400 to estimate φ5300 is 0.36 m2 s−2 . The depth of the shallowest instrument, 5,300 dbar, was chosen as the level to which we referenced the pressure gauges (Donohue et al., 2010) and subsequently φ. This method has been utilized with success in the Gulf Stream (Watts and Johns, 1982), the North Atlantic Current (Meinen and Watts, 2000), Subantarctic Front south of Tasmania (Watts et al., 2001), the Gulf of Mexico (Donohue et al., 2006), and the East China Sea (Andres et al., 2008). Deep pressure and current measurements from the CPIES along with 5,000 m velocities from the seven deep moorings were used to generate optimally interpolated (OI) semi-daily maps of velocity and streamfunction anomaly at 5,300 dbar on 0.1◦ grid. OI mapping is an objective analysis scheme used for data interpolation and mapping (Bretherton et al., 1976). The KESS suite of τ1400 and p time series determined the spatial correlation scales used in OI mapping (not shown). The empirical correlation length scale for τ was 130 km (75 km) for temporal scales greater (less) than 60 days, as determined in Donohue et al. (2010). A two-step approach was used when OI mapping τ following the procedures of

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Tracey et al. (1997), who discussed the mapping method and demonstrated that it sharpens lateral gradients mapped in the Gulf Stream thermocline and improved geostrophic velocity estimates. For deep pressures, after removing an array-wide common mode, a one-step mapping was used with a correlation length scale of 100 km. When comparing the relative strengths of deep- and upper-ocean fluctuations it is necessary that they have the same units. In those cases, bottom pressure was converted to deep streamfunction via ψD = p/ρf , f and ρ are respectively the Coriolis parameter and average deep density. Using the relationship established in Figure 2, τ was transformed into φ5300 , which was subsequently converted to upper-ocean streamfunction ψU = φ5300 /f . For statistical analyses, spectral estimates were calculated using the Welch (1967) averaged periodogram method with 186-day Hanning windows, with 50% overlap. The windowing provided 10.67 equivalent degrees of freedom (EDOF), as estimated for a Hanning N window by EDOF = 83 M , where N is the record length and M is the windowing half-width (Emery and Thompson, 1997). The EDOF for squared coherence is four times larger since we average over four adjacent frequency bands in the ≈ 30–60 day period range. Prior to computing individual power spectra, but not cross-spectra, time series of ψD and ψU were 100-day high-pass filtered in order to remove the more slowly varying field. 4. Observations of deep and upper streamfunction Much of the deep variance is found in the 30–60 day time band, as shown by the spectra of deep streamfunction ψD in Figure 3. Deep energy in the 30–60 day band accounts for 50% of the total deep-pressure variance near the center of the array, while explaining 20% away from the Kuroshio Extension. For reference, 8×107 m4 s−2 streamfunction variance is approximately equivalent to a 9 cm standard deviation in total sea surface height. Sea surface height variance is defined as 2 2 2 σSSH = σψ f 2 /g 2 + σψ f 2 /g 2 + 2CψU ψD f 2 /g 2 , U D

where g is the local gravitational acceleration and C is the covariance (Hendry et al., 2002 and Baker-Yeboah et al., 2009). The deep energy is highest near the center of the array, near the axis of the Kuroshio Extension6 , and decays away from it. For the upper ocean, the spectra ψU are also peaked in the 30–60 day band, with nearly 10 times the variance found in ψD (Fig. 3, middle panel). The highest upper-ocean variance corresponds to 30 cm sea surface height variability. The upper-ocean spectra exhibit additional energy peaks at higher frequencies, which are due to eastward-propagating meanders. The largest variance is along the axis of the Kuroshio Extension with the maximum variance located near the semi-permanent meander trough at 147◦ E. Note that the energy in the 30– 60 day band accounts for 25% of the upper-ocean variance near the meandering Kuroshio. 6. The Kurohsio Extension axis is defined as a mapped τ contour that is equivalent to the Mizuno and White (1983) definition, intersection of the 14◦ isotherm and 300 m depth, and was converted using hydrographic lookup tables in Donohue et al. (2010).

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Figure 3. Left Panel: Variance-preserving spectra of deep-streamfunction anomaly. Middle Panel: Variance-preserving spectra of upper-streamfunction anomaly. In the left and middle panels the line color represents instrument position as indicated by dot color on the right hand panel. Dashed lines indicate the mean position of the Kuroshio Extension jet axis along with its northern and southern boundaries, as determined by the CPIES array. The scale for ψ2U is 10 times that for ψ2D .

Although 75% of observed upper-ocean variance is outside the 30–60 day band, within this period band the upper and deep flow fields exhibit the highest array-wide coherence. Coherence between upper and deep streamfunction measures the local (statistical) vertical coupling in the ocean. Figure 4 shows the squared coherence (left panel) and phase (right panel) between the upper and deep streamfunction, after averaging cross-spectral densities at 0.0161, 0.0215, 0.0269 and 0.0323 cpd frequencies, i.e., over the band of periods ≈ 30–60 days. The 99% confidence level CI for the squared coherence was .10, where the confidence level for the squared coherence was determined using the formula CI = 1 − α1/(EDOF−1) , where α is the % confidence, i.e., 0.01 for 99% or 0.05 for 95% (Emery and Thompson, 1997). Statistically significant coherence between the upper and deep ocean is found throughout the KESS array, with largest values near the axis of the Kuroshio Extension. This large squared coherence suggests that the increased variance in both upper and deep streamfunction in the 30–60 day band is a result of dynamical coupling. Near the Kuroshio Extension, where coherence and energy are highest, the deep fields lead the upper (Fig. 4, right panel). In the region of highest coherence phase lags are ≈90◦ , which corresponds to 10 days. An example of interaction between a deep-pressure center and an upper-ocean meander is shown in Figure 5. On Oct. 14, 2004 (day 287) a deep anticyclone enters from the

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Figure 4. Squared coherence (left panel) and phase (right panel) between deep- and upper-ocean streamfunctions. Cross-spectra were averaged in 0.0161–0.0323 cpd frequency bands (≈ 30–60 day periods). Mean axis of the Kuroshio Extension is denoted by the dashed black line. Coherence and phase is masked for grid points greater than 70 km away from an operating instrument.

northeast corner of the array and approaches an upper-ocean meander crest located at 145◦ E, 36.3◦ N. Initially the anticyclone has pressure anomaly 0.06 dbar and the meander has curvature |κ| = 0.015 km−1 . Between Oct. 22 and Nov. 3, 2004 (days 295–307) the anticyclone encounters the Kuroshio Extension and the meander crest steepens. During the local intensification process the pressure anomaly of the anticyclone and |κ| of the crest increases respectively to 0.14 dbar and 0.033 km−1 . After the interaction the deep anticyclone proceeds SSW, and the meander crest propagates eastward. Figure 6, middle row, demonstrates how frequently deep cyclones and anticyclones propagate SSW across the KESS array. Here, the entire two years of data is used because mapping along the array centerline line is good over the two-year period. It is evident that the large event described in Figure 5 is just one of a recurrent suite of strong events near year days 275–360 and other time intervals. From a wave perspective, there is SSW phase propagation of the individual crests and troughs, but indiscernible SSW group propagation as the packet remains nearly stationary. There are other strong events near days 175–275 and 550–600. Superimposing 30–60 day band-pass filtered upper-ocean streamfunction on the Hovmöller plot (Fig. 6, bottom row) reveals the degree of coupling between the upper and deep fields. When there is strong local intensification of SSW propagating deep-pressure

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Figure 5. An example of a deep anticyclone passing southwestward under a Kuroshio Extension meander crest. Deep streamfunction is color-shaded where anticyclones (cyclones) are represented by the orange (blue) colors. Superimposed in black lines is the upper-ocean streamfunction, with the Kuroshio Extension axis denoted by a white contour line. Four-day snapshots are shown from left to right. As the anticyclone encounters and interacts with the Kuroshio Extension the meander crest steepens.

anomalies, strong upper-ocean anomalies always occur on this same transect. The bottom row of Figure 6 also shows that the deep-ocean streamfunction field leads the upper-ocean streamfunction field in time, consistent with phase lag estimates from the cross-spectra of ψU and ψD (Fig. 4, right panel). Complex empirical orthogonal functions (CEOFs) of 30–60 day band-pass filtered deepand upper-ocean streamfunction fields were utilized to examine the spatial structure of the phase propagation. The first CEOF mode of deep-ocean streamfunction accounts for 64% of the total variance in the 30–60 day band. The amplitude and phase of the first mode for deep streamfunction show SSW phase propagation with maximum spatial amplitude near the Kuroshio Extension meander trough (Fig. 7, top panels).

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Figure 6. Hovmöller diagrams of 30–60 day band-pass filtered deep-ocean streamfunction (middle row), with superimposed 30–60 day band-pass filtered upper-ocean streamfunction (bottom row). For reference, the range of deep-ocean streamfunction values shown correspond to ±11 dbar. Cyclonic (anticyclonic) upper-ocean streamfunction anomalies are denoted by the thin grey (black) contour lines. The Hovmöller section is along a NE-SW line connecting the northern and southernmost CPIES (top row). The white line in the bottom row represents the daily intersection-latitude of the Kuroshio Extension and Hovmöller line section. Gaps or large jumps signify times when the Kuroshio Extension’s path is “S” shaped and has multiple intersection points. (Motivated by personal communication with Hogg and Waterman).

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Figure 7. First mode CEOF for 30–60 day band-pass filtered deep-ocean streamfunction. Mean Kuroshio Extension axis is represented by the dashed black line. Amplitude and phase are masked for grid points greater than 70 km away from an operating instrument. Upper left panel: Spatial amplitude. Upper right panel: Phase. Lower panel: Amplitude time series corresponding to first mode spatial amplitude. The modulus is denoted by the black line and the absolute value of the real part of the amplitude shown in blue. Phase propagation in the direction of increasing phase. This first mode contains 64% of the variance, which is more than twice the next mode (not studied here), and so is distinct from higher modes.

Notably, the phase pattern suggests a plane-wave disturbance, with constant phase lines separated nearly evenly in space. The phase speed calculated from the CEOF is 12–14 km d−1 , which agrees with estimates derived from daily tracking of the deep-pressure anomalies. Wavelengths range from 350–700 km, exceeding 900 km within ≈100 km north of the Kuroshio Extension.

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The phase lines have a wider separation distance, corresponding to faster phase speeds, near the axis of the Kuroshio Extension, suggesting a jet influence on propagation of deep-pressure anomalies. The CEOF amplitude time series has a maximum near yearday 320 (Fig. 7, bottom panel). Comparison with Figure 6 reveals this date coincides with the time interval of multiple strong interactions between SSW propagating deep-pressure anomalies and the Kuroshio Extension. Therefore, the first mode CEOF is capturing the variance structure, spatial and temporal, of SSW propagating deep-pressure anomalies seen in Figures 5 and 6. Correspondingly for the upper ocean, the first mode CEOF of 30–60 day band-pass filtered ψU is shown in Figure 8. Spatial amplitude is large along the Kuroshio axis with southeast phase propagation along the jet in the region of high amplitude. The first mode describes 51% of the total variance in the 30–60 day band. Like the CEOF amplitude time series for the deep ocean, the upper ocean also has its peak near yearday 320. The striking similarity in spatial amplitude patterns and in timing of the peaks of the CEOFs motivated the utilization of a joint CEOF study of upper- and deep-ocean streamfunctions. The two streamfunctions were normalized by the array-averaged standard deviation of each respective field prior to computing the CEOF. By linking the amplitudes of upper and deep ocean in time, a joint CEOF can infer how the upper/deep system behaves as a whole. Figure 9, top left panel, shows deep- and upper-ocean spatial amplitude functions for the joint CEOF. The maximum in the deep-ocean spatial amplitude is located slightly downstream, eastward along the Kuroshio Extension, with respect to the upper-ocean spatial maximum. The joint CEOF upper and deep phase patterns look similar to their respective individual patterns in Figures 7 and 8, so we only show their difference here (Fig. 9, top right panel). Positive phase difference implies that the deep streamfunction leads the upper-ocean streamfunction in time, typically 60–120◦ . The patterns in the CEOF amplitude and phase from the first year of data (Figs. 7–9, top panels) are also present if the full two-year data set is used (not shown). However, in order to calculate CEOFs over the two-year period, OI mapped deep- and upper-ocean streamfunction fields have to be used to fill data gaps due to instrument failure and replacement. We chose not to use OI mapped products in the presented CEOF analysis because it induces site-to-site correlation which could influence the CEOF products. 5. Interpretation and discussion a. Topographically controlled deep-pressure anomalies In this section we pursue a physical model that can describe the observed features of the incoming deep-pressure anomaly centers. The phase θ of the first mode CEOF of deep streamfunction gives the impression of nearly uniform phase propagation, much like a plane wave (Fig. 7, right panel). Given the large wavelengths (350–700 km) and periods (30–60 days) of the signals, and given that estimates that the relative vorticities (P /ρR 2 f ) are usually O(10%f ), we first seek to characterize their kinematics relative to linear TRW theory (Hogg and Waterman, personal communication). Our analyses will apply TRW theory to the point that elements of our observations are found incompatible with TRWs.

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Figure 8. Same as Figure 7 except for 30–60 day band-pass filtered upper-ocean streamfunction. The first mode contains 51% of the variance, which is more than twice the next mode (not studied here), and so is distinct from higher modes.

In the limit that deep stratification is weak, the approximated form of the TRW dispersion relation is given by ω=

f0 ∂H −βk + k Hf00 ∂H ∂y − l H0 ∂x

k2 + l2

,

(1)

where ω is angular frequency, H0 is the mean depth, H is the depth, β is the meridional gradient of f , and k and l are zonal and meridional wavenumbers, respectively. Equation 1 can be re-arranged as γ 2 (β − α)2 + γ2 β−α 2 , (2) + l+ = k+ 2ω 2ω 4ω2

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Figure 9. First-mode joint CEOF of 30–60 day band-pass filtered deep- and upper-ocean streamfunction. Amplitude and phase are masked for grid points greater than 70 km away from an operating instrument. Upper left panel: Spatial amplitude of the deep streamfunction portion of the first mode joint CEOF is shown by color shaded contours. The upper-ocean spatial amplitude is represented by black contour lines. Upper right panel: Phase difference between deep and upper ocean. The mean Kuroshio Extension pathway is denoted by the dashed black line. Lower panel: Amplitude time series of the first joint mode. The first mode contains 54% of the variance, which is more than twice the next mode (not studied here), and so it is distinct from higher modes.

γ f0 ∂H α where α = Hf00 ∂H ∂y and γ = H0 ∂x . For the KESS region β = −.5 and β = .75, having employed version 8.2 bathymetry from Smith and Sandwell (1997), smoothed with a 200 km low-pass filter, to evaluate α and γ. Equation 2 describes constant ω circles in (k, l) space, called “slowness circles.” Figure 10, top row, shows the TRW slowness circles for

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Figure 10. Top row: Barotropic topographic Rossby wave dispersion curves for 30 and 60 days. The grey-shaded region represents the wavenumber space consistent with 30–60 day topographic Rossby waves in the KESS region. Superimposed are the wavenumbers and standard deviation bars associated with the horizontal gradient of phase of the first mode CEOF of 30–60 day band-pass filtered deep streamfunction. Wavenumbers were averaged in 1.5◦ boxes centered on the position indicated by the correspondingly colored dots in the bottom left panel. Bottom right panel shows the number of points included in each 1.5◦ box average. The mean Kuroshio Extension axis is represented with a dashed line in the bottom panels. (Motivated by personal communication with Hogg and Waterman).

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30 and 60 days. This range of periods matches the filtering of the previous section used for calculating the CEOFs. Using the phase of the first mode CEOF of ψD we can estimate the wavenumber, where ∂θ ∂x ∂θ l= . ∂y

k=

(3)

To obtain smoothed wavenumber estimates at locations with a strong signal, the 360◦ phase jumps were first unwrapped and the θ field was smoothed with a 250 km filter prior to taking the gradient. Wavenumbers were only estimated at grid points where the CEOF amplitude exceeded 0.4. In addition, wavenumbers were averaged within 1.5◦ boxes. The average wavenumbers and standard deviation for k and l are superimposed on the TRW dispersion curves (Fig. 10, top panel). Points that lie between the 60-day and 30-day slowness curves, shaded-grey region, are consistent with the TRW dispersion for 30–60 day waves. Most of the smoothed wavenumber estimates do fit a TRW dispersion relation. The wavenumber vector represented by the gray cross inside the 30-day dispersion circle centered on 35◦ N, 147◦ E does not agree with TRW dispersion. It is noteworthy that in this region the deep streamfunction field exhibited its highest variance; it is also the region of highest coherence (strong vertical coupling) between the upper and deep ocean. When incoming pressure anomalies encounter the Kuroshio Extension and couple dynamically via a transient baroclinic growth process, they are no longer free waves and hence may acquire different wavenumber characteristics during the interaction. Our observations show that deep-pressure centers enter from the northeast and translate intact southward and southwestward through the array retaining their approximate size and shape, i.e., without apparent dispersion. We hypothesize that they are nonlinear eddies rather than waves. If the features we observed in the 30-60 day band conformed to purely linear TRW characteristics, with (k, l) as in Figure 10, they would have had a generally north-northwestward group velocity, and the high and low pressure centers would disperse. Thus linear TRW theory is not sufficient to account for the observed eddies, and no energy peak to the southeast is evident (Fig. 7) that might serve as the source. Analogies might first be sought from Chelton et al. (2011), who found that seemingly planetary wave-like disturbances in satellite observations of sea-level anomalies are sometimes more appropriately categorized as nonlinear eddies. As noted by Chelton et al. (2011) and Cushman-Roisin et al. (1990), the translation speed of nonlinear eddies is the long Rossby-wave speed, βRd2 , where Rd is the internal deformation radius. A measure of nonlinearity is the ratio, u/c, of water-parcel speed swirling around the feature to phase speed, and Chelton et al. (2011) found that poleward of 20o the westward phase speed decreased below the commonly observed swirl speeds. This nonlinearity accounted for the translation speeds observed in their analysis of sea-level anomalies.

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Applying these concepts to the Kuroshio Extension, Rd ≈ 40 km, and accounting for the NW-to-SE topographic slope in this region, which effectively causes topographic |β | = 2β, the corresponding TRW long wave speed, 6 cm/s, is less than the observations. Hence, the observed deep eddy translation is not matched by linear topographic wave dispersion and group velocity estimates, nor do estimates for nonlinear long-wave speeds adequately account for their motion. The difference may arise because these eddies may be interacting with the Kuroshio throughout the KESS region, and might not be considered to be freely translating. Just as the non-linear eddies in Chelton et al. (2011) translated along lines of nearly constant f , Kuroshio Extension deep-pressure anomalies travel along f/H lines. Figure 11 shows the phase of the first mode CEOF of 30–60 day bandpass filtered bottom pressures with superimposed f/H contours. As in previous calculations, the topography was smoothed with a lowpass filter. The gradient of the phase is nearly parallel with f/H contours demonstrating the strong effect the regional wide topography has on the propagation direction of the deep-pressure anomalies. A major distinction between eddies and waves arises when u/c > 1 and eddies can carry a core as they translate. Therefore eddies and their energy translate together, in contrast to linear planetary-scale waves for which phase and group velocities can differ in magnitude and direction. The observed features have swirl speeds that cannot be considered small relative to their translation speeds. Even as the deep-pressure anomalies enter from the NNE, prior to interacting with the Kuroshio Extension, their swirl speeds are O(10 cm/s). As they interact with the Kuroshio Extension, they spin up, roughly doubling their pressure anomaly to 0.2 dbar and swirl speed to 20 cm/s. Thus leading up to, and during, the interaction we must consider the ratio u/c = O(1), and the deep-pressure anomalies fall into a gray zone that is neither purely linear waves nor nonlinear eddies. We do find the observations to share many similarities with TRWs, yet their non-dispersive nature is hypothesized to come from their essential non-linear qualities. b. Vertical structure The deep currents associated with the deep-pressure centers are nearly depth-independent. Deep stratification in the Northwestern Pacific is weak with mean subthermocline (2,000 to 6,000 m) buoyancy frequency N = 6 × 10−4 s−1 . The vertical trapping scale Γ ≈ f0 λ/2πN = 8–9 km for a 400 km wavelength wave, and therefore, these deep-eddies should only be weakly bottom-intensified according to linear quasi-geostrophic theory. There were also tall current meter moorings that spanned the Kuroshio Extension measuring currents at 2,000, 3,500 and 5,000 m. The vertical structure of horizontal velocity was approximately depth-independent during strong events. A theoretical trapping scale of 8–9 km is consistent with the findings of Bishop et al. (2012) who used the current velocity magnitudes from the same set of current moorings, along with SSH to examine vertical velocity structure. These separate and comparable estimates of Γ also encouraged our initial comparison of the observed wavenumbers with barotropic TRW dispersion.

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Figure 11. First mode CEOF phase for 30–60 day band-pass filtered deep-ocean streamfunction is shown in shaded color contours. The thin lines denote lines of constant f/H with higher (lower) values in grey (black), i.e., topographic “westward” is towards the southwest.

c. Conceptual model Two-years of observations along the array centerline showed the SSW translation of 12- to 16-deep high and low pressure anomalies across the KESS array (Fig. 6). They arrived in about four sets, and might not be independent of each other. In each instance, the deep-pressure anomalies and upper-ocean meanders intensified, and for simplicity, their individual interactions with the Kuroshio jet will be treated. The evidence from coherence and CEOF studies (Figs. 4 and 9) demonstrated the upper and deep fields have strong statistical coupling. This subsection presents a conceptual model of their dynamical coupling.

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Let us first recapitulate the essential aspects of the observations we seek to explain by reviewing the representative case study shown in Figure 5. A deep-pressure anomaly propagated into the KESS array from the north-northeast. When it approached the baroclinic Kuroshio Extension jet, the interaction affected both the upper and deep fields. For an approaching high, a crest and trough on the upper jet developed and steepened respectively upstream (westward) and downstream (eastward) of the eddy center, while simultaneously the deep-pressure anomaly intensified. For other approaching lows (not shown), a meander trough-crest developed and steepened, but with reversed upstream-downstream sense. The Hovmöller plots (Fig. 6) showed the high and low deep-pressure anomalies led the upperocean anomalies by approximately a one-fourth wave period. This phase shift, interpreted as a spatial phase lag, provides a clue to the dynamical interaction, because that physical arrangement is optimal for growth via baroclinic instability processes. For instance, the fastest growing Eady (1949) mode has the deep streamfunction field leading the upper by ≈ one-fourth wavelength, similar to our setup. However, our process differs from the Eady problem in that our perturbations do not start as infinitesimal, rather they are finite-amplitude deep-pressure centers encountering meanders. As mentioned in the Introduction, Shay et al. (1995) also observed strong vertical coupling events in the Gulf Stream in which meanders and deep-pressure anomalies intensified jointly. However, in their case-studies of joint-development, the initial conditions were different than those in KESS. In SYNOP the deep-pressure anomalies grew from small amplitude pressure features located directly under or near the Gulf Stream, whereas in KESS the deep-pressure anomalies arrive with non-infinitesimal amplitude from an external region. A conceptual model of atmospheric “self development” outlined by Sutcliffe (1947) provides the basic dynamical framework for processes observed in KESS and SYNOP. A simple two-layer ocean model describes the physics of the process. Figure 12 shows a two-layer system with the upper-layer baroclinic jet flowing into the figure plane (lower left panel). As a deep anticyclone approaches a Kuroshio Extension meander crest (top left panel), its nearly depth-independent flow advects water parcels northward in the upper layer within the jet segment upstream of the anticyclone (right hand panels). Note, the term barotropic here refers to the streamfunction at 5,300 dbar, and this streamfunction is used for referencing the upper-layer baroclinic structure. The barotropic component can produce flow normal to the upper baroclinic structure. Since the thermocline shoals northward, in the positive cross-stream direction (+ y dir.), as upper-layer water parcels move across the front they are squashed, and to conserve potential vorticity (PV) they develop negative relative vorticity upstream of the anticyclone. The vorticity generation in turn steepens the crest. The resulting northward excursion of the crest in turn locally squashes the lower layer and to conserve PV, water parcels in the lower layer develop negative relative vorticity, strengthening the deep anticyclone. Hence, the anomaly in each layer acts to strengthen the anomaly in the other layer when the deep anomaly is downstream of an upper anomaly of same sign by approximately a quarter wavelength. This is the essence of joint development as described by Sutcliffe (1947). The corresponding argument with opposite sign can be

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Figure 12. Idealized interaction between an upper-ocean jet and a deep eddy in a two-layer system. The top row is a plane view of a deep anticyclone represented by dashed lines and the upper jet is indicated by solid black lines. The gray dashed line represents the cross-section depicted in the bottom rows. The bottom row shows cross-sections of the jet/anticyclone configuration, with the jet and anticyclone represented by the solid lines and tall grey shaded cylinder, respectively. Left panel: Initial configuration of an eastward-flowing (into the page) jet with an anticyclone downstream of an upper-ocean meander crest. Right panel: Configuration after the crest has steepened and locally the jet has shifted northward (positive y) as indicated by the grey line and the anticyclone has intensified due to the local squashing of the lower water column.

made for deep cyclones approaching the baroclinic jet and developing a meander trough just upstream while jointly intensifying the deep low-pressure anomaly. It is important to distinguish the observational results from baroclinic instability models of Eady (1949) and Phillips (1954). Those linear instability models determined the wavenumber and vertical phase offset of the infinitesimal perturbation that would maximize the exponential growth rate. However, in KESS observations the deep perturbations approach the Kuroshio Extension as finite amplitude deep-pressure anomalies. Furthermore in the models of Eady (1949) and Phillips (1954), upper and deep anomalies propagate with the same along-stream phase speed, ensuring constant lateral phase offset between upper and deep anomalies, maintaining a constant growth rate. In contrast, the observations revealed in the region of strong interaction, 35◦ N, 146◦ E, deep anomalies propagate SSW and upper anomalies propagate SSE. Therefore the structure and growth rate of Eady (1949)

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and Phillips (1954) modes may not be an accurate representation of the observed growth process. d. Numerical process model A numerical process model will next show that the jet need not have pre-existing meanders, and the joint development process described in the previous subsection can also occur when finite amplitude deep-pressure anomalies encounter an initially straight jet. The purpose of the model was to see if southward propagating deep features could couple to an eastward-flowing jet. The dispersion characteristics derived from a CEOF of 30–60 day band-pass filtered deep streamfunction suggested that observed deep-pressure anomalies shared similarities with barotropic TRWs, particularly their phase propagation. There is also evidence that the deep-pressure anomalies came from an external origin, outside the KESS array, and propagated in from the north-northeast (e.g., Fig. 5). The interaction between the deeppressure anomalies and Kuroshio Extension jet generated upper-ocean variability in the 30–60 day band through strong vertical coupling. The coupling process in turn locally modified the dispersion characteristics of the deep-pressure anomalies. Therefore, it is interesting to study the interaction of southward-propagating TRWs with an upper-ocean baroclinic jet over variable bottom slope as an initial-value problem. We chose to simulate TRWs instead of individual deep-pressure anomalies because it is far simpler to model an initially periodic bottom pressure field with periodic boundary conditions. It is beyond the scope of this paper to investigate the effects of all the parameters listed in the next paragraph because our purpose is just to qualitatively illustrate the vertical coupling process. For numerical simulations of the TRWs and an upper-ocean jet we will use the twolayer intermediate geostrophic model described by Sutyrin et al. (2003). The initial setup includes an upper-layer jet without meanders plus periodic barotropic eddies over sloping topography (Fig. 13, top row). When the jet is absent, the barotropic eddies propagate steadily as topographic Rossby waves with the pressure anomaly field described by p sin(kx + kβ t/(k 2 + m2 )) cos(my ), where x = x cos α − y sin α, tan α =

f0 Hx , f0 Hy − βH0

y = x sin α + y cos α,

β = [(f0 Hy /H0 − β)2 + (f0 Hx /H0 )2 ]1/2 .

(4)

The coordinate system has been rotated such that x is along the direction of propagation k, where (k, m) is the wavenumber vector, H0 is the full water column depth, and Hx , Hy are zonal and meridional bottom slope. Choosing L = k −1 as the spatial scale and the inverse wave maximum vorticity T = ζ−1 as the time scale, the flow evolution depends on

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Figure 13. Upper and deep streamfunction fields from numerical model. Black bold lines denote the upper-ocean jet. Solid (dashed) gray contour lines represent the anticyclonic (cyclonic) portion of the topographic wave field. The contour interval for the deep field is .025 dbar. The top and bottom rows show the interaction at elapsed time t = 0 and 30 days, respectively. Initially the upper ocean jet was straight and during the interaction with barotropic TRWs the jet developed steep meander crests and troughs and the TRW amplitude intensified.

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nondimensional values related to the following physical parameters: 1) the effective betaeffect β LT ; 2) the angle of propagation α; 3) the wave aspect ratio mL; 4) the jet velocity Q = V T /L; 5) the radius of deformation γ = Rd /L; 6) depth ratio δ = H1 /(H0 − H1 ), where H1 is the upper layer depth; and 7) Eckman number for bottom friction E = hE /H0 . A simulation for L = 100 km, T = 5 day, Hx = 0.001, Hy = 0, β = 0, m = k, V = .75 m/s, Rd = 50 km, H1 = 1 km, H0 = 6 km, hE = 5 m is shown in Figure 13. For this set of parameters the intrinsic propagation direction, x , is in the negative y direction. In KESS, the zonal topographic slope was 50% larger than the meridional slope, we therefore isolate this effect in the simulation, letting Hy = 0, so that TRWs cross the Kuroshio. The rest of the parameters chosen are representative of the Kuroshio Extension. The idealized case of isolating the physics of the full-water column Kurohsio Extension to an interaction between southward propagating TRWs and an eastward-flowing upper ocean jet can qualitatively reproduce the strong joint development process that steepened meanders and intensified the deep lows and highs in KESS observations. Initially, the pressure amplitude of the TRWs is 0.1 dbar. After 30 days, steep crests and troughs have developed along the jet, during which the deep-pressure anomaly has increased to 0.20 (−0.25) dbar for the anticyclonic (cyclonic) portion of TRW field. As in the conceptual model, the crests and troughs are located where the barotropic flow field of the deeppressure anomalies causes northward (southward) cross-frontal motion. The steepening of the crests and troughs has occurred where the associated squashing or stretching of the lower layer has intensified the deep high- or low-pressure centers respectively. This along-stream phase offset, ≈ one-fourth wavelength, between upper- and deep-pressure anomalies is also in accord with the conceptual model and this enables the joint development. In figure 10 the wavenumber represented by the grey cross was inconsistent with 30–60 day TRWs and we suggested that a vertical coupling process altered phase propagation of pure TRWs near the jet. The numerical process model showed how the interaction between the TRWs and the jet slightly turns the TRW phase speed in the eastward direction. This downstream turning of the phase speed could account for the grey (k,l) pair in Figure 10.

e. OFES model In this section we compare the output from the Ocean General Circulation Model for the Earth Simulator (OFES) against KESS measurements, and in particular we look for evidence of SSW propagating deep-pressure anomalies. Bottom-pressure anomaly was not a variable supplied by the model, but using the provided total sea surface height and vertical temperature and salinity profiles, we calculated bottom-pressure anomaly, as follows: Total sea surface height η can be decomposed into steric ηbc (density varying) and non-steric ηbt (mass loading) components, (Hendry et al., 2002 and Baker-Yeboah et al., 2009) i.e., η = ηbc + ηbt =

φp p + , g ρg

(5)

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Figure 14. Same as figure 6 except 30–60 day band-pass filtered deep- and upper-ocean streamfunction from OFES are included (upper panels), with KESS observations on the bottom panels. All Hovmöller plots have a year base of 2004 and the aspect ratio of each has been adjusted such that an equally sloped line on each represents the same phase speed.

where, φ is referenced to the same level as the pressure anomaly. We chose a reference level of 5,013 dbar because it maximizes the baroclinic signal and is shallow enough to retain many grid points that would otherwise be lost due to topography if we had chosen a deeper level. The 5,013 dbar reference level is similar to that used in KESS, 5,300 dbar. Daily profiles of temperature and salinity were used to construct maps of geopotential anomaly φ5013 (x, y, t), which was then converted into ηbc . Using Equation 5, we calculated daily maps of ηbt and accordingly bottom-pressure anomaly. The pressure anomalies ranged from −0.23 to 0.22 dbar in OFES, consistent with KESS observations −0.28 to 0.27 dbar. Figure 14 shows Hovmöller plots of 30–60 day band-pass filtered deep streamfunction from the OFES model (top panels) and the KESS observations (bottom panels). The transect follows the same line of CPIES in Figure 6. Like the original transect, it was chosen for three reasons: 1) phase lines from the CEOF of bottom streamfunction (Fig. 7, top right panel) are nearly perpendicular to the section; 2) it is the longest line of bottom-pressure measurements in KESS; and 3) it is along a ridge of high variance that extends northeast to

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southwest in the KESS array as seen from the first mode CEOF amplitude. We concentrated on time interval in OFES, Sept. 2001–May 2006, in which strong deep-pressure anomalies propagated SSW through the KESS region. Although not coincident in time, OFES bottom pressures in the 30–60 day band exhibit energetic events such as those seen in KESS. In the ≈ five-year segment shown, the model shows ≈ 50 strong events in which deep-pressure anomalies translated SSW. Like KESS, the upper and deep ocean in OFES are similarly coupled. All cases of strong upper-ocean anomalies accompany a deep-pressure anomaly. The OFES model suggests that this process recurs in packets frequently. 6. Conclusions While meanders and deep-pressure anomalies could develop locally from baroclinic instability, the steepest meanders and most of the highly-energetic deep-pressure anomalies developed when remotely-generated deep-pressure anomalies entered the KESS region from the east and northeast. Their interaction with Kuroshio Extension triggered finite amplitude baroclinic instability processes, which are the subject of this study. The deep energy in the 30–60 day band is strongest near the Kuroshio Extension axis, where deeppressure anomalies intensify, and accounts for as much as 50% of the subtidal deep variance. Daily maps of bottom streamfunction as well as a CEOF of 30–60 day bandpass filtered deep streamfunction revealed that these signals were associated with the SSW propagation of deep-pressure anomalies, where that mode accounted for 64% of variability in that time band. Analysis revealed that the dynamics of the deep-pressure anomalies are neither satisfactorily described by linear TRW theory nor a non-linear eddy description like Chelton et al. (2011). It was hypothesized that the lack of a clear demarcation between linearity and nonlinearity, as gauged by u/c ≈ 1, could be responsible. Nevertheless, the deep-pressure centers translated approximately along f/H contours, consistent with elements of wave theory and with an “eddy” description of the phenomena. The SSW propagating deep-pressure anomalies coupled to the upper Kuroshio Extension jet as they passed under it. Evidence of strong vertical coupling could be seen in the large values of squared coherence between the upper and deep streamfunction fields near the Kuroshio jet axis and in joint CEOFs in the 30–60 day time band. Energy is peaked within this period band and accounts for 25% of the total upper-ocean variability near the axis of the Kuroshio Extension. The numerical simulation demonstrated that the vertical coupling process and enhanced variance near the jet observed in KESS could be simulated by a simple idealization of southward-propagating TRWs interacting with an initially straight eastward-flowing jet. When the deep-pressure anomalies encounter the baroclinic jet, they no longer propagate as free unforced TRWs, but interact with the jet and with neighboring deep-pressure anomalies. This may help explain why in the region of strongest squared coherence in KESS, the CEOF-derived wave numbers deviated from a pattern that was generally in agreement with topographic waves.

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SSW propagating deep cyclones and anticyclones were produced in the general circulation model OFES. The five-year time series revealed that SSW propagation of strong deep-pressure anomalies from an external origin occurs repeatedly in the Kuroshio Extension. Like the KESS observations, the deep OFES anomalies produce a similar range of anomalies in the 30–60 day band and they couple similarly to anticyclonic and cyclonic features in the upper ocean. Although the combined spatial and temporal coverage of the KESS array is unprecedented in Kuroshio Extension, the full mapping capability only existed for one year with continuous time series (June 1, 2004–June 8, 2005) in which to perform these analyses. The five-year time series from OFES, because they strongly resemble the KESS observations, suggest that interactions between non-locally generated deep-pressure anomalies and the Kuroshio Extension is a recurrent process in the ocean. Thus far, we have only remarked on similarity between OFES and KESS in regards to the range of bottom-pressure anomalies and the SSW propagation in the KESS region. In a companion paper we will perform a systematic inter-comparison between KESS observations and OFES, especially in the 30–60 day period band, to validate the veracity of OFES. If proven realistic beyond the metrics we used in this manuscript, OFES could be used to address questions such as: Where and by what process are the incoming deep-pressure anomalies generated and are there other favored regions where deep-pressure anomalies interact strongly with the Kuroshio Extension? Acknowledgments. We are grateful to the U.S. National Science Foundation (NSF) for supporting this under NSF grant OCE 02-21008 and OCE 08-51246. We would like to graciously acknowledge Nelson Hogg, Kathleen Donohue and Stephanie Waterman for their many suggestions involving the scientific interpretation and description of the data. We would also like to thank Jae-Hun Park, Mark Wimbush and Karen Tracey for discussions and processing/management of the data. We wish to thank Gerry Chaplin and Erran Sousa and the URI Equipment Development Lab for preparing and deploying the CPIES array. REFERENCES Andres, M; M. Wimbush; J. H. Park; K. I. Chang; B. H. Lim; D. R. Watts; H. Ichikawa and W. J. Teague. 2008. Observations of Kuroshio flow variations in the East China Sea. J. Geophys. Res., 113, C05013, doi:10.1029/2007JC004200. Baker-Yeboah, S.; D. R. Watts and D. A. Byrne. 2009. Measurements of sea surface height variability in the eastern South Atlantic from pressure-sensor equipped inverted echo sounders: baroclinic and barotropic components. J. Atmos. Oceanic Technol., doi:10.1175/2009JTECHO6959.1 Bishop, Stuart P.; D. Randolph Watts; Jae-Hun Park and Nelson G. Hogg. 2012. Evidence of bottomtrapped currents in the Kuroshio Extension Region.J. Phys. Oceanogr., doi:10.1175/JPO-D-110144.1 Bretherton, F. P.; R. E. Davis and C. B. Fandry. 1976. A technique for objective analysis and design of oceanographic experiments applied to MODE-73. Deep Sea Res., 23, 559–582. Bracco, A. and J. Pedlosky. 2003. Vortex generation by topography in locally unstable baroclinic flows. J. Phys. Oceanogr., 33, 207–219. Chelton, D. B.; M. G. Schlax and R. M Samelson. 2011. Global observations of nonlinear mesoscale eddies. Prog. Oceanogr., 91, 167–216.

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