We perform two additional simulations in which individual causes of O2 variability are successively removed, allowing us to isolate and characterize the contribution of ventilation, circulation, and biology to total O2 changes. We begin by considering AOU along an isopycnal surface, which can be written as the sum of a “preformed” value, AOUo, defined just below the mixed layer, and the total O2 consumption since the water left contact with the mixed layer,
where OUR is the O2 utilization rate, which is integrated along each displacement, ds, of the water parcel as it is transported with velocity, v. The second term on the right-hand side therefore integrates the spatially and temporally varying O2 consumption along the path of the circulation, whose direction and speed are also spatially and temporally variable.
 Preformed AOU is defined at each grid point where an isopycnal layer intersects the mixed layer. This is because while mixed layer O2 is fixed at its saturated value (AOU = 0), the AOU on an isopycnal surface lying just below the mixed layer depends on the degree of exchange with (i.e., ventilation by) the mixed layer. If the detrainment of waters from the winter mixed layer onto an isopycnal surface is large, AOUo will be brought close to zero. If the mixed layer is relatively isolated from the underlying isopycnal surface, the preformed AOU could be much larger than zero.
 From equation (1) we can see that changes in AOU can be caused by changes in its preformed value (AOUo), changes in OUR, or changes in the speed (v) or pathway (ds) of circulation. These contributions to total AOU change can be written schematically as
The first term on the right-hand side (ΔAOUo) represents a change in preformed AOU and is therefore associated with changes in the transfer of O2-rich waters across the base of the mixed layer. We refer to this term as a “ventilation AOU change” (ΔAOUvent). The second term represents an AOU anomaly caused by a change in the circulation (Δcirc), which transports water masses along altered paths (or with altered rates) through the climatological OUR field. We refer to this term as the “circulation AOU change” (ΔAOUcirc) since it is due to the direct affect of changes in water mass location and transport. The last term represents the integrated AOU anomaly due to a change in the distribution of OUR. We refer to this as a “biological AOU change” (ΔAOUbio). It is worth noting that changes in OUR can be caused by both changes in export flux and by changes in the depth of an isopycnal surface.
 Individual sources of AOU variability can now be quantified by performing model integrations in which various terms in equation (2) are eliminated. In order to remove biological AOU changes, we perform a second model integration with the same variable circulation, but this time using the climatological (monthly varying) OUR field diagnosed from the equilibrium spin-up. Holding the pattern of OUR constant through time on each isopycnal surface removes the affect of changes in biological O2 consumption, leaving only ventilation and the direct circulation effects as factors in the variation of O2/AOU.
 In a third simulation, we remove the influence of both ventilation and biological changes on O2/AOU values by holding both OUR and AOUo constant at their climatological values. Changes in AOUo are removed by forcing the AOU on each isopycnal surface to remain at its monthly climatological value wherever that surface intersects the mixed layer either in the climatological mean state or at any time during the variable circulation run. For example, the region in which a given isopycnal surface intersects the mixed layer could shift poleward as a result of warming. This would produce a new region of isopycnal/mixed layer interaction to the north of the climatological outcrop. The effect of increased ventilation on O2/AOU in that region is eliminated in this simulation, by forcing AOU at the local isopycnal/mixed layer interface to remain at climatological values. Similarly, equatorward of the climatological outcrop the isopycnal layer will lose contact with the mixed layer, as less dense water intercedes between them. The impact on O2 of the resulting decrease in ventilation is again prevented, by maintaining the AOU at its climatological value at that location. This manipulation removes the effect of changes in the location and/or strength of ventilation, assuring that changes in the detrainment of mixed layer water onto an isopycnal surface will have no effect on the preformed AOU. Holding constant both OUR and AOUo therefore yields AOU changes that result solely from the direct influence of circulation changes.
 Table 1 summarizes the three simulations and the sources of AOU variability in each one. The contribution of biology, ventilation, and circulation to the total AOU change can thus be separated as follows:
The sum of AOU anomalies from these three processes yields the total AOU change. Finally, we note that because the ventilation and circulation of the thermocline are dynamically coupled processes it is impossible to strictly separate their impacts on O2 distributions. What we seek instead is an approximation of the direct role of each process, which must be evaluated a posteriori for consistency with basic features of ocean circulation (e.g., ventilation-induced O2 changes should originate in regions where ventilation occurs) and related patterns of model variability (e.g., biologically induced O2 changes should resemble patterns of export flux anomalies).
Table 1. Summary of Simulations Used to Isolate Components of Model AOU Variabilitya
|Experiment||Description||AOU Terms Present|
|1||full variability||ΔAOUbio + ΔAOUvent + ΔAOUcirc|
|2||climatological OUR field||ΔAOUvent + ΔAOUcirc|
|3||climatological OUR field, and climatological preformed AOU||ΔAOUcirc|
 The contributions of biology, circulation, and ventilation changes to the simulated O2 difference between the 1980s and the 1990s are shown in Figures 8 and 9. Each map represents the contribution to the decadal O2 difference of a single process integrated along the path of circulation since the beginning of the simulation. For example, the ventilation O2 change between the 1980s and 1990s includes not only the local effect at the outcrop location of ventilation changes occurring in those 2 decades, but the affect of ventilation changes from 1948 to 1990 throughout the basin. For this reason, a ventilation O2 change may be observed far from the isopycnal outcrop, corresponding to changes in the more distant past that have been transported and dispersed by the circulation. The sum of the O2 changes attributed to all three processes, together with the small change in O2sat, is equal to the total O2 difference over the specified period.
Figure 8. Difference between decadal mean O2 (μmol/kg) in the 1990s and the 1980s along the isopycnal surface σθ 26.6. The (a) total O2 difference is the sum of O2 anomalies due to (b) biological changes, (c) ventilation changes, (d) changes in circulation, and thermodynamically driven O2 changes, which are negligible (not shown). The decadal O2 difference is nearly equal to but opposite of the total decadal AOU change.
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 On σθ 26.6, O2 changes are predominantly caused by physical processes of circulation and ventilation. The direct impact of changes in circulation constitutes the largest source of O2 changes in the tropical-to-subtropical latitudes, while ventilation is an important factor in much of the subpolar region. In the western subpolar gyre, ventilation changes are the primary cause of O2 decreases. Decreased O2 in the northeastern Pacific is caused roughly equally by changes in ventilation and circulation. The effect of reduced ventilation may still be underestimated, however, because the model does not account for changes in surface fresh water forcing. The salinity of North Pacific Intermediate Water has been shown to have decreased during recent decades [Wong et al., 1999a]. The freshening of surface waters in this region would increase the stratification of the upper water column, likely amplifying the reduction in ventilation and the associated AOU increase in the model's subpolar region. The large band of increased O2 extending across the subtropical gyre is almost entirely circulation driven. O2 changes attributable to biology are very small (<5 μM), except in the tropics and in a small region of the Subarctic North Pacific, where biological anomalies are an important part of the total O2 change. Changes in O2sat are insignificant (≪5 μmol/kg) everywhere except in a narrow region off the east coast of Japan, where temperature gradients are large.
 In the central mode water (σθ 25.8, Figure 9) the response of O2 to ventilation changes is spatially complex and includes both O2 decreases along the Kuroshio Extension and positive anomalies across much of the rest of the northern basin. These ventilation-induced O2 changes reflect changes in the ventilation rate of Central Mode Water that are positive in some regions and negative in others [Ladd and Thompson, 2002]. Circulation related O2 changes are confined to the tropics, where O2 concentrations decrease from the 1980s to the 1990s to the south of 10°N, but increase north of 10°N. Biologically induced O2 changes are more important on this density surface, where decadal differences in excess of 10 μM occur in several regions. Decreased O2 in an eastern tropical band along 10°N and in the subtropical-subpolar transition zone (30°N–40°N) correspond to regions of increased export from the 1980s to the 1990s, while reduced export along ∼15°N allows O2 to increase in underlying waters (Figure 10).
Figure 10. Ventilation rates from 1948 to 2000 on isopycnal layers (a) σθ 25.4 representing subtropical mode water, (b) σθ 25.8 and (c) σθ 26.2 representing central mode water, and (d) the base of the ventilated thermocline (σθ 26.6). Rates are given as a mean detrainment velocity of waters from the winter mixed layer onto the isopycnal layer.
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 Because most of the remineralization of sinking organic matter occurs above 300 m, the impact of biological variability on O2 is concentrated in the shallow water column. The strong decrease in sinking flux with depth also means that the signature of export changes on thermocline O2 is spread across many isopycnal surfaces. Export anomalies in the center of the subtropical gyre, where isopycnals are deep, are seen primarily on light density surfaces, whereas a tropical export change affects O2 on denser surfaces as well (see Figures 8b and 9b). For this reason, O2 changes observed along constant density surfaces bear only a weak resemblance to surface productivity changes, even when productivity changes are significant. When biological O2 anomalies are plotted on constant depth surfaces (not shown), they are seen to closely match the pattern of export changes.
 While biologically driven O2 changes on σθ 25.8 are significant, they are offset to a large degree by physically driven anomalies. For example, the widespread biological O2 decreases in the central North Pacific are roughly balanced in that region by increased O2 due to circulation and ventilation. The close coupling of biologically and physically driven O2 variability is due to the biological compensation mechanism discussed above, in which an increase in O2 supply to the ocean interior is accompanied by an overall increase in nutrient supply back to the surface layer. If such changes in nutrient supply are not balanced by changes in organic matter export, as assumed in our model, this coupling could be weaker than suggested by our results. Nevertheless, unless organic matter export is completely decoupled from changes in nutrient supply, O2 trends that arise from changes in the biological pump are likely to be difficult to detect owing to counteracting physically driven O2 changes. A better understanding of historical changes in surface nutrient concentrations would help to constrain the strength of biological-physical compensation in the variability of thermocline O2.
4.2. Origins of ΔAOUvent and ΔAOUcirc
 The attribution of simulated O2 changes indicates that ventilation and circulation are the dominant drivers of lower thermocline O2 variability in the late twentieth century. Here we trace the origins of the primary O2 anomalies to specific changes in lower thermocline ventilation and circulation. We find that O2 differences between the 1980s and 1990s are due to both decadal trends in ventilation and circulation, and to episodic physical perturbations that generate large-scale propagating O2 anomalies with decadal lifetimes.
 Ventilation rates among different isopycnal layers respond differently to atmospheric forcing (Figure 10). All layers, however, exhibit significant interannual variability over the course of the hindcast simulation. On σθ 26.6, ventilation rates decrease throughout the simulation, owing largely to a reduction in the area of the wintertime outcrop. This reduces the flux of well-oxygenated waters onto this isopycnal, causing O2 decreases from the 1970s through the 1990s (Figure 5). The reduced outcrop area of denser isopycnals is compensated by an increase in the area of outcropping among shallower isopycnals, whose wintertime outcrops expand northward. The increase in outcrop area, together with intensified wind speeds in the northwest Pacific leads to enhanced overall ventilation rates from the 1970s to the 1990s (Figure 10a) among subtropical mode waters (σθ < 25.8).
 While decadal changes in ventilation rate account for much of the O2 decrease in the subarctic Pacific below σθ 26.0, decadal circulation changes account for much of the subtropical O2 increases. The origin of this simulated decadal O2 change is a southward expansion of the subtropical gyre from the 1980s to the 1990s, causing low-O2 waters in the North Equatorial Countercurrent to be displaced along ∼15°N–20°N by relatively well oxygenated waters in the gyre's North Equatorial Current lying just to the north. The observed O2 changes along 152°W south of 25°N (Figure 4), together with concurrent salinity changes [Emerson et al., 2001], support such an expansion of the subtropical gyre.
 Circulation related O2 changes are also observed in the Kuroshio/Oyashio Extension (KOE) region, where strong meridional O2 gradients are produced by the confluence of warm tropical/subtropical waters of the Kuroshio current and cold subpolar waters of the Oyashio current. The position of the model KOE oscillates narrowly between a more northern and southern location. Fluctuations in both the strength and position of the Kuroshio Extension produce decadal O2 deviations throughout the simulation (Figure 11a). Some shifts occur along the entire current, causing simultaneous O2 changes from the western boundary to the central Pacific (e.g., early 1970s, Figure 11a). Other shifts, for example in the late 1970s, are focused in the western boundary, where the O2 anomalies are generated and transported eastward (Figure 11a). The timing and western intensification of this shift in the Kuroshio Extension, coincident with the well-known shift in the Pacific Decadal Oscillation, are features that have been observed in historical hydrographic data [Deser et al., 1999].
Figure 11. Time series of O2 concentrations (minus the mean value during the 1960s) at locations in the subtropical North Pacific showing the propagation of anomalies (a, b) from the Kuroshio/Oyashio Extension and (d) from the eastern tropical North Pacific into (c, e) the subtropical gyre. In each case, the O2 anomaly originates in the early 1970s, but remains an identifiable feature into the 1990s. (f) Locations of the time series given in Figures 11a–11e.
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 Decadal changes in circulation and ventilation explain much of the O2 anomaly pattern, but shorter-term episodic events also contribute substantially to O2 variability. Some physical perturbations generate O2 anomalies of sufficient scale and intensity that they remain coherent features for over a decade. Such anomalies are capable of traveling far from their regions of origin, becoming progressively broader but lower in amplitude as they are transported across the basin (Figure 11c).
 For example, the multidecadal trend toward reduced ventilation of σθ 26.6 is punctuated by a brief period of increased O2-rich mass flux in the early 1970s (Figure 10d), when a new region of ventilation opens in the central Pacific along 35°N–40°N. This ventilation event injects a large bolus of high O2 water into the permanent thermocline (Figure 11b), adding to the elevated O2 originating in the KOE. The high O2 anomaly continues eastward and then southward into the subtropical gyre, producing first an increase in O2 along 150° W from the 1970s to the 1980s (Figure 5) followed by a decrease from the 1980s to the 1990s. This transient response to the temporary increase in ventilation on σθ 26.6 contributes to the simulated O2 decrease coincident with that observed by Emerson et al. .
 Episodic physical perturbations also contribute to the model's subtropical O2 increases. For example, the O2 increase in the center of the gyre on σθ 26.6 is the remnant signal of a transient O2 anomaly transported westward from the eastern tropical North Pacific (ETNP). In the early 1970s, low-O2 waters carried north by the coastal California Undercurrent are briefly entrained into the broad, southward flowing California Current. The resulting mass of O2-depleted water is carried into the subtropical gyre by the California and North Equatorial Currents (Figures 11d and 11e), causing first local O2 decreases (e.g., west of Hawaii from the 1970s to the 1980s, Figure 7) followed by local O2 increases (e.g., same region from the 80s to 90s, Figure 6) as the anomaly is transported through the region and O2 levels return to their background levels. Evidence for the presence of low -O2 water masses originating in the ETNP has been observed near Hawaii by [Lukas and Santiago-Mandujano, 2001].
 To summarize, the differences in O2 between the 1980s and 1990s at the base of the ventilated thermocline are caused by a combination of decadal trends in circulation and ventilation, and transient responses to brief, localized physical perturbations (Figure 12 and Table 2). In the subpolar region, O2 decreases on σθ 26.6 are due to a multidecadal reduction in lower thermocline ventilation rates. In the subtropics, O2 increases at the southeastern gyre boundary are due to the southern expansion of the gyre from the 1980s to the 1990s. At the subtropical/subpolar transition zone, a large positive O2 anomaly is generated in the early 1970s by a displacement of the Kuroshio Extension and the brief onset of ventilation in the central Pacific. At the same time, a bolus of low-O2 water is entrained into the California Current from the oxygen minimum zone in the east. Both of these anomalous water masses are subsequently transported into the gyre, causing regional trends in the 1970s to the 1980s that reverse sign into the 1990s as the anomaly is dissipated and transported out of the region.
Figure 12. Schematic summary of the origins of major O2 changes from the 1980s to the 1990s in the lower ventilated thermocline. The physical causes of anomalies labeled a through d are identified in Table 2 and described in the text.
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Table 2. Specific Physical Causes of Simulated Decadal O2 Changes Between the 1980s and 1990s on Isopycnal σθ 26.6a
|ΔO2 (1990s–1980s) on σθ 26.6||Ventilation||Circulation|
|Decadal trends||subpolar O2 decrease, Figure 12, label a||subtropical O2 increase, Figure 12, label d|
|Interannual perturbations||subtropical O2 decrease (preceded by an increase, 1970s to 1980s), Figure 12, label b||subtropical O2 increase (preceded by a decrease, 1970s to 1980s), Figure 12, label c|
4.3. Changes in O2 Inventory
 The mean O2 anomaly averaged over the entire basin thermocline from the base of the mixed layer to σθ 27.0 (Figure 13) reveals an interannual variability of ∼5 μM, but no long-term change during the course of the simulation. Thus the O2 inventory of the North Pacific remains relatively stable over the period 1950–2000 despite the presence of significant regional trends. Using the separation of O2 changes due to biological, ventilation and circulation, we can quantify the effect of each process on O2 inventory.
Figure 13. Changes in average O2 concentrations relative to the 1960s, for the entire thermocline (σθ < 27.0) of the model North Pacific basin. Mean O2 anomalies (solid black line) are decomposed into contributions from changes in O2 solubility (solid gray line), ventilation (dashed black line), circulation (dotted black line), biology (dashed gray line) according to equation (2). Biologically and physically driven O2 changes act to maintain a constant O2 inventory over the course of the simulation, although interannual variations of ∼5 μM are evident.
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 The response of the mean basin O2 concentration due to variability in physical processes of circulation and ventilation is an increase of roughly 4 μM (Figure 13). The simulated expansion of the subtropical gyre into formerly suboxic waters of the ETNP accounts for at least part of the circulation-driven O2 increase. The net effect of variable ventilation rates also enhances the supply of O2, when averaged across the basin. Increased ventilation rates among shallower surfaces evidently dominate the reduced ventilation of the lower ventilated thermocline (σθ 26.6), the process responsible for significant subpolar O2 decreases. Why do we not see large O2 increases in the shallow thermocline associated with the increased ventilation rates?
 The intensification of North Pacific circulation causes an increasing supply of nutrients to the surface and therefore an increase in subsurface biological O2 consumption. The resulting biological O2 anomalies averaged over the entire basin decrease with time and closely compensate the combined anomalies due to ventilation and circulation, thus stabilizing the total O2 burden on the decadal timescale. The coupling of biological and physical O2 changes is seen regionally on σθ 25.8 (Figure 9). In the central North Pacific, regions of enhanced ventilation are often coincident with increased biological O2 consumption due to elevated export flux. In the absence of changes in biological productivity, physically driven O2 changes would therefore be observed throughout the water column. Instead, physically driven changes are largely compensated by biologically driven O2 changes in the shallow water column, where oxygen utilization rates are relatively high. In contrast, the lower ventilated thermocline experiences relatively low rates of O2 consumption, and physically driven O2 variability remains largely uncompensated.
 The close coupling of biologically and physically driven O2 variability is in part a consequence of maintaining constant surface nutrients. In the absence of a change in surface nutrients, it is sea surface temperatures (i.e., the thermodynamic boundary condition for O2 saturation; see Figure 13) that govern the O2 inventory. In the absence of changes in surface nutrients and SST, the O2 content of the ocean interior cannot change for more than a few years.