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Research article 31 Jul 2019
Research article | 31 Jul 2019
Correspondence: Jian Liu (njdllj@126.com) and Zhengyu Liu (liu.7022@osu.edu)
HideCorrespondence: Jian Liu (njdllj@126.com) and Zhengyu Liu (liu.7022@osu.edu)
Previous studies show that the evolution of global mean temperature forced by the total forcing is almost the same as the sum of individual orbital, ice sheet, greenhouse gas and meltwater single forcing runs in the last 12 000 years in three independent climate models: Community Climate System Model 3 (CCSM3), Fast Met Office/UK Universities Simulator (FAMOUS) and Loch-Vecode-Ecbilt-Clio-Agism Model (LOVECLIM). This validity of the linear response is useful because it simplifies the interpretation of the climate evolution. However, it has remained unclear if this linear response is valid on other spatial and temporal scales and, if valid, in what regions. Here, using a set of TraCE-21ka (Simulation of the Transient Climate of the Last 21,000 Years) climate simulations, the spatial and temporal dependence of the linear response of the surface temperature evolution in the Holocene is assessed approximately using the correlation coefficient and a linear error index. The results show that the response of global mean temperature is almost linear on orbital, millennial and centennial scales in the Holocene but not on a decadal scale. The linear response differs significantly between the Northern Hemisphere (NH) and Southern Hemisphere (SH). In the NH, the response is almost linear on a millennial scale, while in the SH the response is almost linear on an orbital scale. Furthermore, at regional scales, the linear responses differ substantially between the orbital, millennial, centennial and decadal timescales. On an orbital scale, the linear response is dominant for most regions, even in a small area of a midsize country like Germany. On a millennial scale, the response is still approximately linear in the NH over many regions. Relatively, the linear response is degenerated somewhat over most regions in the SH. On the centennial and decadal timescales, the response is no longer linear in almost all the regions. The regions where the response is linear on the millennial scale are mostly consistent with those on the orbital scale, notably western Eurasian, North Africa, subtropical North Pacific, the tropical Atlantic and the Indian Ocean, likely causing a large signal-to-noise ratio over these regions. This finding will be helpful for improving our understanding of the regional climate response to various climate forcing factors in the Holocene, especially on orbital and millennial scales.
Long-term temperature evolution in the Pleistocene is often believed to be, and therefore interpreted as being, driven mainly by several external forcing factors, notably, orbital forcing, greenhouse gases (GHGs), continental ice sheets and meltwater flux forcing. (Here, we treat the coupled ocean–atmosphere system as our climate system, such that Earth orbital parameters, GHGs, meltwater discharge and continental ice sheet are regarded as external forcing.) Implicit in this interpretation is often an assumption that the response is almost linear to the four forcing factors, that is, the temperature evolution forced by the total forcing combined is approximately the same as the sum of the temperature responses forced individually by the four forcing factors. This linear response, if valid, simplifies the interpretation of the climate evolution dramatically because each feature of the climate evolution can now be attributed to those on different forcing factors. One example is the global mean temperature evolution of the last 21 000 years (COHMAP members, 1988; Liu et al., 2014). It has been shown that the global mean temperature response is almost linear to the four forcing factors above in three independent climate models (Community Climate System Model 3 (CCSM3), Fast Met Office/UK Universities Simulator (FAMOUS) and Loch-Vecode-Ecbilt-Clio-Agism Model (LOVECLIM); Fig. 2 of Liu et al., 2014) with the temperature evolution forced by the total forcing almost the same as the sum of those individually forced by each forcing factor. Furthermore, this deglacial warming response is forced predominantly by the increase in GHGs, with significant contribution from the ice sheet retreat. This linear response, however, has not been assessed quantitatively for the climate evolution in the Holocene. The Holocene period poses a more stringent and interesting test of the linear response, as it removes the deglacial global warming response that is dominated by that to increased CO_{2} and ice sheet retreating (Fig. 2a of Liu et al., 2014). An even more interesting and practical question here is, besides the global mean temperature response to the slow orbital timescale, how linear is the response at shorter temporal scales and smaller spatial scales, throughout the Holocene?
In general, the assessment of the linear response, in principle, can be done in a climate model using a set of experiments that are forced by the combined forcing as well as each individual forcing. Furthermore, each forcing experiment has to consist of a large number of ensemble members. This follows because a single realization of a coupled ocean–atmosphere model could contain strong internal climate variability on a wide range of timescales (Laepple and Huybers, 2014), from daily variability of synoptic weather storms (Hasselmann, 1976) to interannual variability of El Niño (Cobb et al., 2013) and interdecadal climate variability (Delworth and Mann, 2000), all the way to millennial climate variability (Bond et al., 1997). The ensemble mean is therefore necessary to suppress internal variability and then generate the truly forced response to each forcing. The goodness of the linear response can therefore be assessed by comparing the response to the total forcing with the sum of the individual responses. One practical problem with this ensemble approach is, however, the extraordinary computing costs, especially for long experiments in more realistic, fully coupled general circulation models. A more practical question is therefore: is it possible to obtain a meaningful assessment of the linear response using only a single realization of each forced experiment for the Holocene, such as those in TraCE-21ka (Simulation of the Transient Climate of the Last 21,000 Years) experiments (Liu et al., 2014).
Strictly speaking, it is impossible to disentangle the forced response from internal variability in a single realization. This would make the assessment of the linear response difficult. However, it is conceivable that, if our interest is the slow climate evolution of millennial or longer timescales in response to the slow forcing factors such as the orbital forcing, ice sheet forcing, GHGs and meltwater flux, the assessment is still possible, albeit approximately, at least for very large-scale variability. This follows because these forcing factors are of long timescales and of large spatial scales; the forced response signal should therefore also be on long timescales and large spatial scales if the response is approximately linear. An extreme example is the almost linear response in the global temperature of the last 21 000 years as discussed by Liu et al. (2014). In contrast, internal variability in the coupled ocean–atmosphere system tends to be of shorter timescales (decadal to centennial) and of smaller spatial scales, at least in the current generation of coupled ocean–atmosphere models. This naturally leads to two questions. First, how linear is the climate response at different spatial and temporal scales, quantitatively? Second, in what regions does the linear response tend to dominate? The answer to these questions should help improve our understanding of regional climate response during the Holocene. A further question is as follows: if the linear approximation is valid, what is the contribution of each forcing factor in different regions and at different timescales. This question will be addressed in a follow-up paper (Wan et al., 2019).
In this paper, we assess the linear response for the Holocene temperature evolution quantitatively, using five forced climate simulations in CCSM3 (Liu et al., 2014), with the focus on the spatial and temporal dependence of the linear response. We will assess the linearity response to orbital, millennial, centennial and decadal timescales and on global, hemispheric and regional spatial scales. The data and methodology are given in Sect. 2. The dependence of the linear response to spatial and temporal scales is analyzed in Sect. 3. A summary and further discussions are given in Sect. 4.
The data are from TraCE-21ka (Liu et al., 2009, 2014), which consists of a set of five synchronously coupled atmosphere-ocean general circulation model simulations for the last 21 000 years. The simulations are completed using the CCSM3 (Community Climate System Model version 3). The simulation forced by the total forcing (experiment ALL) is forced by realistic continental ice sheets, the GHGs, orbital forcing and meltwater fluxes. The ice sheet is changed approximately once every 500 years, according to the ICE-5G reconstruction (Peltier, 2004). The atmospheric GHG concentration is derived from the reconstruction of Joos and Spahni (2008). The orbital forcing follows that of Berger (1978). The coastlines at the LGM (Last Glacial Maximum) were also taken from the ICE-5G reconstruction and were modified at 13.1, 12.9, 7.6 and 6.2 ka, after which the transient simulation adopted the present-day coastlines. The meltwater flux follows largely the reconstructed sea level and other paleoclimate information and, in the meantime, reconciles the response of Greenland temperature and AMOC (Atlantic Meridional Overturning Circulation) strength in comparison with reconstructions. More information on the details of the experiment and forcing can be seen in He (2011) or on the TraCE-21ka website http://www.cgd.ucar.edu/ccr/TraCE/ (last access: 11 July 2019).
The transient simulation under the total climate forcing reproduces many large-scale features of the deglacial climate evolution consistent with the observations (Shakun et al., 2012; Marsicek et al., 2018), suggesting a potentially reasonable climate sensitivity in CCSM3, at global and continental scales. In addition to the all-forcing run (ALL), there are four individual forcing runs forced by the orbital forcing (ORB), the continental ice sheets (ICE), the GHGs (GHG) and meltwater forcing (MWF) (Liu et al., 2014; Table 1). In these four experiments, only one forcing varies the same as in experiment ALL, while other forcings/conditions remain the same as at 19 ka. Therefore, this set of experiments can be used to study the linear response of the climate to the four forcing factors. Here, we will only examine the surface temperature response in the Holocene (last 11 000 years).
We will use correlation and normalized root means square error (RMSE) to assess the linear response (see next subsection for details). We note, however, that our assessment of the linear response is approximate. Before introducing the details of the assessment method, it is useful here to make some general comments on the linear response assessment. As pointed out by one reviewer, strictly speaking, the assessment of the linear response requires one to answer two questions.
How linear is the response to external forcing?
What is the relative importance of external forcing vs. internal variability, assuming the response were linear?
Specifically, for Q1, if we denote the temperature response to the full external forcing by T_{R}(F_{all}(t)), the response to the individual forcings by T_{R}(F_{i}(t)) (with $i=\mathrm{1},\mathrm{\dots},\mathrm{4}$), and the internal temperature variability of the five model simulations by T_{I,all}, T_{I,1}, T_{I,2}, T_{I,3} and T_{I,4}, the linearity of the response could be defined by the extent to which the total forced response equals the sum of the individual responses or
In our case, we only have a single member for each experiment as
and
and the linearity is assessed from the correlation (and normalized RMSE) between the sum of the individual experiments $\sum _{i=\mathrm{1}}^{\mathrm{4}}{T}_{i}\left(t\right)$ and the total forcing experiment T_{all}(t). Therefore, our linearity assessment is contaminated by the noise of internal variability. This can be seen, for example, in the correlation as
Here, 〈 〉 indicates covariance, and we have assumed that the forced response T_{R}(F_{*}(t)) and internal variability ${T}_{\text{I},*}$ (where $*=$ all, 1, 2, 3, 4) are independent of each other; in addition, the time series is sufficiently long, so sampling errors are negligible. The correlation therefore depends on the signal-to-noise ratio. If the noise (internal variability) is large, the correlation will be much smaller than 1 even if the response is purely linear as in Eq. (1). The only way to suppress internal variability is to perform a large number of ensemble simulations for each experiment. Given only one member for each experiment, we have to be content that our linear assessment using Eq. (2) is approximate, depending on the signal-to-noise ratio. Related to this problem of single member experiments, since we cannot distinguish internal variability from forced response clearly, Q2 cannot be assessed exactly either.
In spite of these potential issues, with a single member for each experiment, useful information can still be extracted on linear response. Our general hypothesis is that the slow (orbital and millennial) and large (continental and basin) variability is composed mostly of forced signal and the faster (centennial and shorter) and smaller variability is mostly associated with internal variability of noise. In other words, in our set of single member of simulations, the signal-to-noise ratio is large for slow variability but small for faster variability. Qualitatively, this hypothesis seems reasonable. First, all the four external forcing factors are of slow timescales and large spatial scales; additionally, internal variability is usually weak in the coupled ocean–atmosphere system at slow timescales and large spatial scales. Our focus is indeed the slow variability and large scale here, so we can roughly treat the slow and large variability in the single realization as the signal and the linearity of the response may be assessed using Eq. (2). Second, again, because our forcing factors are of slow timescales and large spatial scales, higher-frequency or small-scale variability in the model should not be dominated by the forced variability (unless the response is highly nonlinear). Therefore, high-frequency or small-scale variability can be treated roughly as “noise”. This is consistent with the later assessment that slow variability seems to be an approximately linear response while high-frequency variability not. Based on this hypothesis, the signal-to-noise ratio is also estimated using the variance of slow variability as the signal and high-frequency variability as the noise (as in late Fig. 7). It should be noted however that this hypothesis is qualitative in nature. One major purpose of this paper is to give a somewhat more quantitative assessment on this hypothesis. How slow, how large and how good will the linear response be?
Our experimental design is proper for linear response assessment here. Alternatively, in another experimental setting, individual forcing experiments are often superimposed sequentially one by one: for example, first the ice sheet, second the ice sheet plus orbital forcing, third the ice sheet, orbital and GHGs, and finally, applying all four forcings of ice sheet, orbital, GHGs and meltwater. In this experimental design, the full forcing response is by default the response of the sum response after adding the four forcing factors together, and therefore it cannot be used to assess the linearity of the response. Nevertheless, it should be kept in mind that our four individual forcing experiments are not designed optimally for the study of the linear response in the Holocene. This is because, except for the variable forcing, all the other three forcing factors are fixed at the 19 ka condition. As such, the mean state is perturbed from the glacial state, instead of a Holocene state. This may have contributed to some unknown deterioration on the linear response discussed later. Nevertheless, we believe, our major conclusion should hold reasonably well. This is because, partly, the response is indeed almost linear for orbital and millennial variability as will be shown later.
We use two indices to evaluate the linear response: the temporal correlation coefficient r and a normalized linear error index L_{e}. The correlation coefficient is calculated as
Here, ${S}_{t}=\sum _{i=\mathrm{1}}^{\mathrm{4}}{T}_{i}/\mathrm{4}$ is the linear sum of the temperature time series T_{i} of the four single forcing experiments, T_{t} is the full temperature time series in the ALL run (both at time t) and n is the length of the time series. The overbar represents the time mean. The correlation coefficient represents the similarity of the temporal evolution between the sum response and the ALL response. However, the correlation does not address the magnitude of the response. Indeed, even if S_{t} and T_{t} has a perfect correlation r=1, the two time series can still differ by an arbitrary constant in their magnitudes. Therefore, we will also use a normalized linear error index L_{e} to evaluate the magnitude of the linear response. Here, L_{e} is defined as the RMSE of the sum temperature response from the full temperature response divided by the standard deviation of the full temperature response in the ALL run:
In general, a large r (close to 1) and a smaller L_{e} (close to 0) represents a better linear approximation, with r=1 and L_{e}=0 as a perfect linear response. Therefore, if r is close to 1 and L_{e} is close to 0, we can conclude that the response is close to linear. As noted above, with a single realization here, our assessment of the linear response has limitations. First, if r is sufficiently small and L_{e} is sufficiently large, we cannot confirm the response is either linear or nonlinear because the small r or large L_{e} can also be contributed by strong internal variability. Second, if the forcing is dominated by shorter timescale variability, say interannual to interdecadal variability, as in the case of volcanic forcing or solar variability, it will be difficult to assess the linear response. This because the timescales of the forced response now overlap heavily with strong internal climate variability in the coupled system, and it will be difficult to separate the forced response from internal variability without an ensemble mean.
But how to assess the goodness of the linear response from the value of r and L_{e}? We can test the goodness of the linear response statistically on r and L_{e}.
The statistical significance of r for a particular timescale is tested using the Monte Carlo method (Kroese, 2011, 2014; Kastner, 2010; Binder, 1997) with 1 000 000 realizations on the corresponding red noise in the AR(1) model (autoregressive model of order 1) which uses the AR(1) coefficient derived from the model to generate time series. The fit uses the lag-1 auto-correlation coefficient.
The statistical significance of the L_{e} of a particular timescale is tested using a bootstrap method (Efron, 1979, 1993) with 1 000 000 realizations on the corresponding time series. Specifically, the bootstrap is done as follows, taking the global mean temperature as example. First, we will derive the L_{e} from one random realization on the temperature of the ALL run of the 100 binned data (110 points of data, each representing a 100-year bin). For this random realization, the order of the original temperature time series is swapped randomly. Then, this realization is used as a new ALL response for comparison with the sum response of the four individual experiments to derive a L_{e} in Eq. (2). Since the random realization distorts the serial correlation time with the sum response, one should usually expect a large error L_{e}. Second, we repeat this process for 1 000 000 times on 1 000 000 random realizations; this will produce 1 000 000 random values of L_{e}, forming the PDF of the L_{e}. Third, the minimum 5 % level is then used as the 95 % confidence level.
The dependence of the linear response on spatial and temporal scales will be studied by filtering the time series on different scales. For the spatial scale, we will divide the globe into nine successive cases, denoted by nine division factors: f=0, 1, 2, 3, 4, 6, 8, 12 and 24 from the largest global scale to the smallest model grid scale. The f=0 case is for the global average, while the f=1 case is for the hemispheric average in the Northern Hemisphere (NH) and Southern Hemisphere (SH). Further division will be done within each hemisphere. Note that each hemisphere has 96 (long.)×24 (lat) grid boxes, with a ratio of 4 : 1 between longitude span and latitude span. We divide each hemisphere into f×f sections of equal latitude and longitude spans, with each area containing the same number of $(\mathrm{96}/f)\times (\mathrm{24}/f)$ grid boxes, maintaining the ratio of 4 : 1 between longitude span and latitude span. For example, f=2 is for the 2×2 division, with each area containing 48×12 grid boxes; f=24 is for the 24×24 division with each area containing 4×1 grid boxes, about the size of $\mathrm{15}{}^{\circ}\left(\text{long.}\right)\times \mathrm{3.75}{}^{\circ}\left(\text{lat}\right)$, like the size of a midsize country such as Spain or Germany in the midlatitudes.
On the timescale, we decompose a full 11 000-year annual temperature time series (from 11 to 0 ka) in 100-year bins (a total of 110 data bins or points, each representing a 100-year mean) into three components. The three components are to represent the variability of, roughly, orbital, millennial and centennial timescales. Following Marsicek et al. (2018), we derive the orbital and millennial variability using a low-pass filter called the locally weighted regression fits (Loess fits) (Cleveland, 1979). First, the orbital variability is derived by applying a 6500-year Loess fit low-pass filter to the temperature time series, and it therefore contains the trend and the slow evolution longer than ∼ 6500 years. Second, we apply a 2500-year Loess fit low-pass filter onto the temperature time series; then, we derive the millennial variability using this 2500-year low-pass data subtracting the 6500-year low-pass data. Finally, centennial variability is derived as the difference between the 100-year binned temperature time series and the 2500-year low-pass time series. In addition, we also derive a decadal variability time series. First, we compile the 10-year bin time series from the original 11 000-year annual time series (of a total of 1100 data points, each representing a 10-year mean). Second, we apply a 100-year running mean low-pass filter to the time series of the 10-year binned data. Finally, decadal variability is derived by using the 10-year binned time series minus its 100-year running mean time series.
Given the different degrees of freedom especially among the filtered variability of different timescales, it is important to test the goodness of the linear response statistically on r and L_{e} on different timescales differently. As a reference, the significance level is tested against the global mean temperature series in the ALL run. For the total, orbital, millennial, centennial and decadal temperature time series, the 95 % confidence levels are found to be 0.72 (with the AR(1) coefficient 0.96), 0.76 (0.97), 0.65 (0.95), 0.21 (0.31) and 0.19 (0.06), respectively.
In this paper, this AR(1) test for global mean temperature is also used as the common significant test for different spatial scales and in different regions as well. This use of a common significance level is for simplicity here. First, the use of different regional AR(1) coefficients for different regions will make the comparison of the linear responses among different spatial scales (e.g., Figs. 3 and 4) and different regions (Figs. 5 and 6) difficult. Second, except for the orbital timescale, the AR(1) coefficient for the global mean temperature is larger than most of the regional AR(1) coefficients (not shown), likely caused by the further suppression of internal variability in the global mean. As a result, the global mean AR(1) test actually serves as a more stringent test than the local AR(1) test. At the orbital scale, the global mean AR(1) coefficient is in about the middle of the regional AR(1) coefficients. The uncertainly of using the global mean AR(1) coefficient is therefore about the average of those of regional AR(1) coefficients. Third, and, most importantly, as our first study here, our focus is on the global features of the linear response. The difference among the AR(1) coefficients among different spatial scales and different regions is much smaller than that between different timescales here. Therefore, the global mean AR(1) can still provide an approximate guideline for the proper significant test at different timescales. In later studies, if one's focus is on a specific spatial scale and on a specific region, the regional AR(1) should be used to reexamine the significance test.
As a reference, the significance level of L_{e} is tested against the global mean temperature series in the ALL run. For the total, orbital, millennial, centennial and decadal temperature time series, the 95 % confidence levels are found to be 1.23, 1.23, 1.21, 1.24 and 1.36, respectively. This suggests that when the RMSE is less than about 1.2–1.3 times of the total response, the linear sum is not significantly different from the total response at the 95 % confidence level. As for the L_{e} test, since our focus here is on the global feature of the linear response, for simplicity, the significance level derived from the global mean temperature is used as the common confidence level for all regional scales.
The global mean temperature provides a useful example to start the discussion of the dependence of the linear response on timescales. We first examine the linear response of the global mean temperature based on its components of orbital, millennial, centennial and decadal variability (Fig. 1). Figure 1a is the total variability of global surface temperature derived from the ALL run and the sum of the four individual forcing experiments. The global temperature response is almost linear on the orbital and millennial scales throughout the Holocene (Fig. 1b and c). The orbital-scale evolution is characterized by a warming trend of about 1^{∘} from 11 ka to ∼ 4.5 ka before decreasing slightly afterwards. This feature is captured in the linear sum, albeit with a slightly smaller magnitude and an additional local minimum around 3 ka (Fig. 1b). The total variability is very similar to the orbital variability (r=0.99, Fig. 1a vs. Fig. 1b). The millennial variability shows five major peaks around ∼ 9.8, 7.8, 4.7, 3.7 and 1.8 ka. All these peaks seem to be captured in the sum response, albeit with a slightly larger amplitude (Fig. 1c). For orbital and millennial variability, the correlation coefficients between the sum and the full responses are r=0.83 and 0.71, respectively, both significant at the 95 % confidence level and explaining over 50 % of the variance; the linear errors are L_{e}=0.63 and 0.92, respectively, also significant at the 95 % confidence level. It should be noted, however, that the goodness of the linear response is based on the entire period and is meant for the response of the timescale to be studied. Therefore, even for a good linear response at long timescales, the sum response may still differ from the total response significant at some particular time. For example, for the orbital-scale response in Fig. 1b, even though the linear response is good in terms of r and L_{e}, there is a 1 ^{∘}C difference between the sum and total responses at 11 and 3 ka. Therefore, for the orbital-scale response, the linear response mainly refers to the trend-like slow response of the comparable timescale of the orbital scale instead of some response features of shorter timescales. Further down the scale, at a centennial timescale, the global centennial variability appears also to exhibit a modest linear response (Fig. 1d), with r=0.44 and L_{e}=1.21. But the linear response of the decadal variability becomes poor (Fig. 1e), with $r=-\mathrm{0.02}$ and L_{e}=1.99, which is not statistically significant at the 95 % confidence level. The result of the analysis of global mean temperature is qualitatively consistent with the previous hypothesis that the linear response tends to degenerate at a shorter temporal scale because of the smaller forced response signal and the presence of strong internal variability. It shows that, for global temperature, the response is approximately linear at orbital and millennial timescales but becomes much less so at centennial scales and fails completely at decadal scales.
In order to assess the linear response at different spatial scales, we first analyze the linear response to the hemisphere scale for the NH and SH (f=1). It is interesting that the linear response significantly differs between the NH and SH (Fig. 2). Figure 2a and f are the total variability of hemispheric surface temperature of the ALL response and sum response for the NH and SH, respectively. Their components on the four timescales (i.e., orbital, millennial, centennial and decadal) are shown in Fig. 2b–e and 2g–j, respectively. In the NH, the response is almost linear at the millennial scale (r=0.82, L_{e}=1.01, Fig. 2c) but not so strong on orbital (r=0.55, L_{e}=0.84, Fig. 2b) and centennial (r=0.32, L_{e}=2.29, Fig. 2d) scales, while only L_{e} is significant on the orbital scale, and only r is significant on the centennial scale. At the decadal scale, the linear response fails completely ($r=-\mathrm{0.04}$, L_{e}=1.99, Fig. 2e). In comparison, in the SH, the linear response is dominant at the orbital scale (r=0.92, L_{e}=0.43, Fig. 2g) but poor on all other timescales, including millennial ($r=-\mathrm{0.12}$, L_{e}=2.32, Fig. 2h), centennial (r=0.14, L_{e}=3.19, Fig. 2i) and decadal (r=0.03, L_{e}=2.07, Fig. 2j) scales. The linear response of the global mean temperature discussed in Fig. 1 therefore seems to be dominated by the SH response to the orbital scale but by the NH response to the millennial scale. This suggests that the goodness of the linear response depends on both the region and timescale, highlighting the need to study the linear response at regional scales.
The linear response at different spatial scales and on the orbital, millennial, centennial and decadal timescales is summarized in Figs. 3 and 4 in the correlation coefficient and linear error index, respectively. Figure 3a shows the correlation coefficients of the orbital variability in each region for the nine division factors. The cases of global mean (f=0) and hemispheric mean (f=1) have been discussed in detail before. The correlation coefficients for succeeding division factors ($f=\mathrm{2},\mathrm{3},\mathrm{\dots},\mathrm{24}$) show several features. First, as expected, the correlation coefficient tends to decrease towards smaller areas (larger f). Quantitatively, however, the correlation coefficient does not decrease much, such that even in the smallest area (f=24), the correlation in most regions is still above 0.8, statistically significant at the 95 % confidence level. This suggests that the response at the orbital scale is almost linear over most regions, even at the smallest scale of about a midsize country like Germany (f=24, 15^{∘}(long.) × 3.75^{∘}(lat)). Second, the linear response in the NH is slightly better than SH (for f≥3), a topic to be returned to later. Third, subareas in both hemispheres show a comparable linear response across all the spatial scales, with the median correlations all above ∼ 0.8, except that of the NH mean temperature (f=1). The linear response of NH is not better than those of regional variability at smaller spatial scales (f≥2). This is opposite to the expectation that the linear response becomes more distinct for a larger area because the average over a larger area tends to suppress internal variability more. This case, however, seems to be a special feature and should be treated with caution. The correlation coefficient therefore shows that, for orbital-scale evolution, the temperature response is dominated by the linear response over most of the globe, even at regional scales. These features are also consistent with the linear error analysis in Fig. 4a.
Millennial variability also shows a weaker linear response for smaller scales, in both the correlation (Fig. 3b) and linear error (Fig. 4b). Quantitatively, for millennial variability, the response is still approximately linear in the NH over many regions, albeit less so than at the orbital scale. The correlation coefficients remain above 0.6 across most regions even at the smallest division area (f=24), contributing to ∼ 40 % of the variance. In contrast to orbital variability, where regions in both hemispheres show a comparable linear response, millennial variability shows that the response cannot be confirmed as linear in most regions in the SH, with the median no longer significant at the 95 % confidence level. Similar to the orbital variability, nevertheless, the responses are more linear in the NH than SH on all the spatial scales.
In contrast to orbital and millennial variability, almost no response can be confirmed as linear for centennial variability. The median linear response is no longer significant on the centennial timescale in either hemisphere across spatial scales (f>3, Figs. 3c and 4c), with few correlation coefficients larger than 0.3 and contributing less than 10 % of the variance. Finally, decadal variability exhibits absolutely no linear response over any spatial scales in either the NH or SH (Figs. 3d and 4d).
The approximate linear response at the orbital and millennial scales suggests that these two groups of variability are generated predominantly by the external forcing. In contrast, the poor linear response of centennial and decadal variability suggests that these two groups of variability are caused mainly by the internal coupled ocean–atmosphere processes. This is largely consistent with our original hypothesis. It should be kept in mind that, in our single realization here, the poor linear response to centennial and decadal variability may also be contributed by nonlinear responses of the climate system. But, given the almost complete absence of forcing variability at this short timescale in our experiments, we do not think that the nonlinear response is the major cause of the poor linear response here.
We now further study the pattern of the linear response. Figure 5 shows the spatial patterns of the correlation coefficients at orbital (Fig. 5a1–a3) and millennial (Fig. 5b1–b3) scales for three representative spatial scales: f=3, 6 and 24 (the other factors are not shown because they are similar to the abovementioned three representative spatial scales). For orbital variability (Fig. 5a1–a3), the response is almost linear in most regions in the NH on all three spatial scales, with the correlation coefficients above 0.8. In the SH, the response is also almost linear over the continents, except for over Australia but is not linear over the ocean. This leads to the significantly reduced linear response in the SH as discussed in Figs. 3a–4a. This suggests that orbital variability is likely forced predominantly by external forcing over continents. The overall poorer linear response over ocean than land, however, is puzzling. The orbital timescale is so long that one would expect a similar quasi-equilibrium response over both land and ocean surface. This issue deserves further study in the future. More specifically, at the regional scales, e.g., f=6 and 24 (Fig. 5a2 and a3), the response is almost linear over the western half of the Eurasian continent, North Africa, Central and South America, most NH oceans and SH tropical oceans, and the Antarctic continent, as seen in Fig. 5a2 and a3 (only significant correlation coefficients are shown). But the linear response is poor over the North America continent and the eastern Eurasian continent, and the entire Southern Ocean. Similar features can also be seen in the map of linear error (not shown).
For millennial variability, in the NH, the linear response shows a similar feature to that of orbital variability, but the linear response is poor over almost all the SH. Figure 5b1–b3 show that the response is almost linear in most regions in the NH at the three spatial scales, with the correlation coefficient above 0.6. At the regional scale, e.g., f=6 and 24 (Fig. 5b2 and b3), the response is almost linear over the northwestern Eurasian continent, North Africa, northern North America, the northern Pacific Ocean, the southern North Atlantic Ocean and the western Arctic Ocean, as seen in Fig. 5b2 and b3 (only significant correlation coefficients are shown). But the linear response is poor over southern North America and the eastern and southern Eurasian continent. While in the SH, the linear relationship is poor over almost the entire SH as seen in the correlation map (Fig. 5b1–b3). Interestingly, over the NH, the regions of the linear response for millennial variability are mostly consistent with the regions of the linear response for orbital variability, notably western Eurasia, North Africa, the subtropical North Pacific, the tropical Atlantic and the Indian Ocean. These preferred region of the linear response potentially suggests some common mechanisms of the climate response in these regions, in this model.
In order to understand the cause of the preferred regions of the linear response, we examine the signal-to-noise ratio. As discussed in Sect. 2.2, our forcing factors are on millennial and orbital timescales, and the linear response is also largely valid for orbital and millennial variability. We will therefore use the variance of the orbital and millennial variability as a crude estimate of the linear response signal. Similarly, since there is no centennial and decadal forcing in our model and the response of centennial and decadal variability is not a linear response, we use the variance of the sum of the centennial and decadal variability as a rough estimate for internal variability as the linear noise. Admittedly, this estimation is crude, limited by the single realization here. This signal-to-noise ratio does not directly address Q2 in Sect. 2.2 because the timescales of the signal and noise are different. Instead, it is used as a rough estimation of the relative magnitude of the signal-to-noise ratio between different regions, with the assumption that the relative noise level between different regions may be not too sensitive to the timescales. Indeed, the use of the signal-to-noise ratio here is to shed some light on the regional preference of the linear response. Figure 6 shows the signal-to-noise ratio for orbital and millennial variability for the three representative spatial scales (f=3, 6 and 24). For orbital variability (Fig. 6a1–a3), the signal-to-noise ratio is large (above 10^{(0.6)}, the log base 10 is taken on the signal-to-noise ratios) in most regions in the NH. In the SH, the signal-to-noise ratio is also large over the continents, but is small over the ocean. At different regional scales, e.g., f=6 and 24 (Fig. 6a2 and a3), the signal-to-noise ratio is large over the western half of the Eurasian continent, North Africa, Central and South America, most NH oceans and SH tropical oceans, and the Antarctic Continent. But the signal-to-noise ratio is small over Canada and the eastern Eurasian continent, and the entire Southern Ocean. These spatial features of signal-to-noise ratio on the orbital scale (Fig. 6a1–a3) are similar to those in the correlation map (Fig. 5a1–a3). The spatial correlation between the map of the signal-to-noise ratio in Fig. 6 and the corresponding correlation coefficient in Fig. 5 is 0.53, 0.58 and 0.49 for f=3, 6 and 24, respectively, as seen in the scatter diagram in Fig. 7, all significant at the 95 % confidence level.
For millennial variability, the signal-to-noise ratio also shows a similar feature to that of orbital variability although overall somewhat smaller (note the different color scales). Figure 6b1–b3 show that the signal-to-noise ratio is large in most regions in the NH in all three spatial scales, with the signal-to-noise ratio above 10^{(−1)} (the log base 10 is taken on the signal-to-noise ratios). At the regional scale, e.g., f=6 and 24 (Fig. 6b2 and b3), the signal-to-noise ratio is large over the northwestern Eurasian continent, North Africa, central North America, the northern North Pacific Ocean, the North Atlantic Ocean and the Arctic Ocean. But the signal-to-noise ratio is small over the North American continent outside the central North America, South America, and the eastern and southern Eurasian continent. While the signal-to-noise ratio is small over almost the entire SH. Over the NH, interestingly, the regions of large signal-to-noise ratio for millennial variability are mostly consistent with the regions of large signal-to-noise ratio for orbital variability, notably northwestern Eurasia, North Africa, the subtropical North Pacific, the North Atlantic and the tropical north Indian Ocean. These features of the spatial pattern of signal-to-noise ratio on the millennial scale (Fig. 6b1–b3) are similar to those in the correlation map (Fig. 5b1–b3). The correlation coefficient between the maps of signal-to-noise ratio and correlation is 0.65, 0.51 and 0.45 for f=3, 6 and 24, respectively (Fig. 7), again all significant at the 95 % confidence level.
In this paper, the linear response is assessed for the surface temperature response to orbital forcing, GHGs, meltwater discharge and continental ice sheet throughout the Holocene in a coupled GCM (general circulation model; CCSM3). The global mean temperature response is almost linear on the orbital, millennial and even centennial scales throughout the Holocene but not for decadal variability (Fig. 1). Furthermore, the sum response accounts for over 50 % of the total response variance for orbital and millennial variability. Further analysis on the regional scale suggests that the response is approximately linear on the orbital and millennial scales for most continental regions over the NH and SH, with the sum response explaining over about 50 % of the total response variance. However, the linear response is not significant over much of the ocean, especially over the ocean in the SH. There are specific regions where the linear response tends to be dominant, notably the western Eurasian continent, North Africa, central and South America, the Antarctic continent, and the North Pacific. The strong linear response is interpreted as the region of large signal-to-noise ratio. That is, in these regions, either the orbital and millennial response signal is large or the influence of the centennial and decadal variability noise is small or both. This suggests that the orbital and millennial variability in these regions is relatively easy to understand. This finding lays a foundation for our further understanding of the impacts of different climate forcing factors on the temperature evolution in the Holocene of orbital and millennial timescales. This understanding is our original motivation for this work. Further work is underway in understanding the contribution of different forcing factors on the temperature evolution (Wan et al., 2019).
It should be kept in mind that since there is only one member for each experiment, we cannot separate the forced response signal from the internal variability of noise clearly at each timescale. Therefore, we cannot address Q1 and Q2 raised in Sect. 2.2 accurately. Instead, our assessment is likely contaminated by internal variability (see discussions in Sects. 1 and 2). In particular, for smaller-scale variability, of which internal variability is likely to be strong and the forced signal is likely to be weak, our correlation may underestimate the linearity of the response (see Eq. 2). Nevertheless, we speculate that our results on large-scale variability still remain robust. Furthermore, at regional scales, although the absolute value of the linear correlation of the forced response may be underestimated, it is possible that relative between different regions, the linear assessment may still be somewhat valid. These speculations, however, require much further study, especially with ensemble experiments. In spite of its limitation, our study represents the first systematic assessment of the linear response for the Holocene and can serve as a starting point for further studies in the future.
There are many further issues that need to be studied. Our study here is carried out for a single variable (surface temperature) in a single model (CCSM3) for the Holocene. Yet, the linear response could differ for different variables, in different models, for different periods and for different sets of forcing factors. For example, if we evaluate the precipitation response in the Holocene in CCSM3, the response is less linear than temperature (not shown); this is expected because the precipitation response contains more internal variability and exhibits more nonlinear behavior than temperature. The assessment will be also different if a different period is assessed, e.g., the last 21 000 years; with a large amplitude of climate forcing, the linear response may degenerate in the 21 000-year period. In addition, the assessment of the linear response using only one realization will be difficult to perform for volcanic forcing and solar variability forcing; these forcing factors have short timescales and therefore their impacts will be difficult to separate from internal variability without ensemble experiments. Finally, it is also important to repeat the same assessment here in different models and to establish the robustness of the assessment. It should also be kept in mind that our assessment is implicitly related to the assumption that, at millennial and orbital timescales, internal variability is not strong relative to the forced responses. Although this seems to be consistent in our model, there is a possibility that internal variability is severely underestimated in the model compared to the real world (Laepple and Huybers, 2014). If true, the relevance of our model assessment to the real world will be limited. It should also be kept in mind that, if the response is dominated by that to a single forcing, the assessment of the linear response here becomes one that is more relevant to the question of the forced response vs. internal variability, as discussed in Q2 in Sect. 2.2. As a further step, though, one can examine if the magnitude of the total response responds to the magnitude of this single forcing linearly.
Even in the context of this model assessment, much further work remains. Most importantly, the purpose of testing the linear response is for a better understanding of the physical mechanism of the climate response. It is highly desirable to understand why the response tends to be linear in some regions but not in others. In particular, it is unclear why the linear response is preferred over land than over ocean for orbital and millennial variability. At such a long timescale, one would expect that the upper-ocean response has reached quasi-equilibrium and therefore the surface temperature response over land and over ocean should not be too different. Ultimately, we would like to assess and understand the physical mechanism of the climate evolution in different regions. This work is underway (Wan et al., 2019).
The TraCE-21ka data sets were provided by NCAR from their website at https://www.earthsystemgrid.org/project/trace.html (last access: 11 July 2019).
LW, JL and ZL conceived the idea for the paper; LW carried out the analysis and prepared the first draft. ZL contributed great ideas and gave the suggestion for the analysis. JL and ZL revised the paper several times. WS provided some useful suggestions to the paper. BL helped to download the larger TraCE-21ka data. All coauthors helped to improve the paper.
The authors declare that they have no conflict of interest.
The authors thank Kai Ding for checking the grammar of the first draft. The TraCE-21ka data sets were provided by NCAR.
This work was jointly supported by the National Key Research and Development Program of China (grant no. 2016YFA0600401), the National Natural Science Foundation of China (grant no. 41420104002, 41630527), the Program of Innovative Research Team of Jiangsu Higher Education Institutions of China and the Priority Academic Program Development of Jiangsu Higher Education Institutions (grant no. 164320H116), and NSF OCN1810681.
This paper was edited by Qiuzhen Yin and reviewed by Oliver Bothe and two anonymous referees.
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